Download to the PDF file. - CURVE

Petrography, Geochemistry and Isotopic Analysis of Paleogene Volcanism in the
Fish Creek Mountains, Great Basin, North-Central Nevada
by
Christopher Stevens
A thesis submitted to the Faculty of Graduate and Postdoctoral
Affairs in Partial Fulfillment o f the requirements
for the degree of
Master of Science
in
Earth Sciences
Carleton University
Ottawa, Ontario
©2013
Christopher Stevens
1+1
Library and Archives
Canada
Bibliotheque et
Archives Canada
Published Heritage
Branch
Direction du
Patrimoine de I'edition
395 Wellington Street
Ottawa ON K1A0N4
Canada
395, rue Wellington
Ottawa ON K1A 0N4
Canada
Your file Votre reference
ISBN:
978-0-494-94308-3
Our file Notre reference
ISBN:
978-0-494-94308-3
NOTICE:
AVIS:
The author has granted a non­
exclusive license allowing Library and
Archives Canada to reproduce,
publish, archive, preserve, conserve,
communicate to the public by
telecommunication or on the Internet,
loan, distrbute and sell theses
worldwide, for commercial or non­
commercial purposes, in microform,
paper, electronic and/or any other
formats.
L'auteur a accorde une licence non exclusive
permettant a la Bibliotheque et Archives
Canada de reproduire, publier, archiver,
sauvegarder, conserver, transmettre au public
par telecommunication ou par I'lnternet, preter,
distribuer et vendre des theses partout dans le
monde, a des fins commerciales ou autres, sur
support microforme, papier, electronique et/ou
autres formats.
The author retains copyright
ownership and moral rights in this
thesis. Neither the thesis nor
substantial extracts from it may be
printed or otherwise reproduced
without the author's permission.
L'auteur conserve la propriete du droit d'auteur
et des droits moraux qui protege cette these. Ni
la these ni des extraits substantiels de celle-ci
ne doivent etre imprimes ou autrement
reproduits sans son autorisation.
In compliance with the Canadian
Privacy Act some supporting forms
may have been removed from this
thesis.
Conform em ent a la loi canadienne sur la
protection de la vie privee, quelques
formulaires secondaires ont ete enleves de
cette these.
W hile these forms may be included
in the document page count, their
removal does not represent any loss
of content from the thesis.
Bien que ces formulaires aient inclus dans
la pagination, il n'y aura aucun contenu
manquant.
Canada
ABSTRACT
The Fish Creek Mountains, located in north-central Nevada, is a site o f multiple igneous
events ranging from 35Ma to IMa, covering most of the igneous history o f the Great
Basin where my goal of this project is to investigate the Paleogene volcanism within the
FCM. Samples collected from the FCM and surrounding Tobin and Shoshone Ranges are
Eocene-Oligocene in age (33-34 Ma), calc-alkaline basaltic andesites through to
rhyolites. Incompatible elements and isotopic data suggest continental margin
subduction-like trace element signatures with highly radiogenic 87Sr/86Sr, consistent with
an old, metasomatized lithospheric mantle source. Compared to other areas o f the
Ancestral Cascades and Western Great Basin, elevated 87Sr/86Sr found in the FCM lavas
are herein interpreted to be the result of a purely lithospheric component, suggesting a
thicker crust in the central Great Basin. These intermediate, mantle-derived magmas
ascended and mixed with partial melts of the mafic crust to form andesites and dacite.
ACKNOWLEDGEMENTS
I would like to thank my supervisor, Brian Cousens for giving me the opportunity
to do this project and for his continual aid, patience and encouragement throughout this
study.
I would also like to thank Dr. Lizzy Ann Spencer and Prof. John Blenkinsop
(members of the Isotope Geochemistry and Geochronology Research Centre) for thenoutstanding patience and expertise where as training myself on all the various
geochemical research tools. Thanks to A1 Alcazar (thin section preparation laboratory)
for his help in preparing my thin sections.
I would like to thank my friends at Carleton who put aside their own work to help
me when I needed it. Finally, special thanks are given to my family who has always
supported me throughout my years studying geology.
TABLE OF CONTENTS
ABSTRACT............................................................................................................................. ii
ACKNOWLEDGEMENTS................................................................................................... iii
Table of Contents....................................................................................................................iv
List of Figures.......................................................................................................................... vi
List of symbols, Nomenclature or Abbreviations................................................................ xi
CHAPTER 1 - INTRODUCTION.......................................................................................... 1
1.1 Introduction............................................................................................................. 1
1.2 Justification or purpose for this study...................................................................3
1.3 Geology and tectonic setting................................................................................. 4
1.3.1 Brief geochronology of the Great Basin........................................... 4
1.3.2 Great Basin igneous assemblages and volcanism............................. 9
1.3.3 Great Basin Tectonics....................................................................... 13
1.4 Local geology (Nevada).......................................................................................15
1.5 Components in continental subduction zones.................................................... 17
1.6 Degree of crustal contamination.........................................................................20
CHAPTER 2 - METHODOLOGY...................................................................................... 22
2.1 Field Methods...................................................................................................... 22
2.2 Petrography.......................................................................................................... 22
2.3 Geochemistry powder preparation......................................................................23
2.4 Major element geochemistry - X-Ray Fluorescence (XRF)............................ 24
2.5 Trace element geochemistry - Inductively Coupled Plasma Mass
Spectrometry (ICP-MS)...................................................................................... 24
2.6 Isotope Geochemistry - Thermal Ionization Mass Spectrometry (TIMS)
26
CHAPTER 3 - GEOLOGY AND FIELD RELATIONSHIPS IN THE FISH CREEK
MOUNTAINS......................................................................................................................... 29
3.1 Physical Observations..........................................................................................29
CHAPTER 4 - GEOCHRONOLOGY................................................................................. 40
4.1 Introduction......................................................................................................... 40
4.2 Previous dating.................................................................................................... 40
4.3 New Ar-Ar ages...................................................................................................42
CHAPTER 5 - PETROGRAPHY OF PALEOGENE VOLCANIC ROCKS OF THE
FCM....................................................................................................................................... 43
5.1 Introduction......................................................................................................... 43
5.2 Basaltic Andesites...............................................................................................43
5.3 Andesites............................................................................................................. 44
5.4 Dacites................................................................................................................. 48
5.5 Rhyolites............................................................................................................. 48
5.6 Summary............................................................................................................. 51
CHAPTER 6 - GEOCHEMISTRY; GEOCHEMICAL AND RADIOGENIC ISOTOPE
SYSTEMATICS OF PALEOGENE FELSIC TO MAFIC ROCKS OF THE FISH
CREEK MOUNTAINS, NORTH-CENTRAL NEVADA, WESTERN UNITED
STATES................................................................................................................................. 54
6.1 Introduction........................................................................................................ 54
6.2 Whole rock major geochemistry.......................................................................54
6.3 Whole rock trace element geochemistry.......................................................... 60
6.4 Isotope Geochemistry........................................................................................ 6 6
CHAPTER 7 - DISCUSSION..............................................................................................70
7.1 Introduction......................................................................................................... 70
7.2 Petrological comparisons.................................................................................... 70
7.3 Geochemical comparisons.................................................................................. 74
7.4 Sources of the Parental M elts............................................................................79
7.5 Dacite and Rhyolite genesis............................................................................... 94
7.6 Andesites............................................................................................................ 99
7.7 Evidence of crustal contamination....................................................................104
7.8 Mantle beneath the FCM and tectonic model.................................................. 108
CONCLUSIONS..................................................................................................................113
REFERENCES.....................................................................................................................116
Appendix A: Rock descriptions and TAS name with sample number and location
124
Appendix B: Thin Section Summary................................................................................. 134
Appendix C: X-ray Fluorescence major element data collected in this study..................136
Appendix D: ICP-MS trace element data collected in this stu d y .................................... 138
Appendix E: Mass Spectrometry isotopic data collected in this study............................ 143
LIST OF FIGURES
Figure 1: Location of Fish Creek Mountains study area and surrounding volcanic areas o f
the western United States. From John (2001).........................................................................2
Figure 2: Position of the Great Basin in the Western Cordillera. Black star represents
study location. From Dickinson (2006).................................................................................. 5
Figure 3: East-dipping continental margin subduction zone with associated periods of
heightened magmatic activity and deformation related to the subduction o f the Farallon
plate under North America. A) Latest Mesozoic-earliest Cenozoic B) Middle to late
Cenozoic. Figures from Fiero (1986) and Zhang et al. (2009).............................................. 8
Figure 4: Maps showing the general distribution o f Cenozoic volcanic assemblages
outlined in dark grey. A) interior andesite-rhyolite assemblage, B) Western andesite
assemblage and C) Bimodal basalt-rhyolite assemblage. Figures modified from John
(2001)..................................................................................................................................... 12
Figure 5: Rock sample locations for samples collected during the 2010 field season
plotted as blue crosses on the topographic map of the FCM and surrounding area. Red
crosses represent 24.9 Ma rhyolitic tuff. Modified from Fish Creek Mountains and
Edwards Creek Valley USGS topographic maps, U.S. Geological Survey, 1:100 000
scale........................................................................................................................................30
Figure 6 : Typical basaltic andesite representative sample (10-BV-39) from the Tobin
Range in outcrop (A) and hand sample (B)..........................................................................33
Figure 7: Sample 10-BV-41 in outcrop (A) and hand sample (B). This andesite was taken
from an outcrop that could have been a collapsed dome or a debris flow that shows
vertical joints, lithic clasts and boulders on top yet has a volcanic matrix filled with
crystals................................................................................................................................... 33
Figure 8 : Sample 10-BV-10 in outcrop (A) and in hand sample (B) characteristic of nonvesicular flows from the western FCM suit. They are mostly porphyritic andesites taken
from columnar jointed outcrops........................................................................................... 34
Figure 9: Sample 10-BV-06 in outcrop (A) and hand sample (B). This is the most
northern expression of Paleogene volcanics in the western FCM suite and is suspected to
have been part o f a lava flow top or a’a ...............................................................................34
Figure 10: Sample 10-BV-45 in outcrop (A) shows crude columnar jointing of fresh lava
and in hand sample contains trace anhedral olivine with minor plagioclase (B) sampled
from the Shoshone Range...................................................................................................... 36
Figure 11: Sample 10-BV-47 in outcrop (A) and hand sample (B) characteristic of
samples taken from platy and fissile outcrops of the Shoshone range............................... 36
Figure 12: Dacitic-rhyolitic cone of sample 10-BV-48 from the Shoshone range which
appeared to have lava tunnels up the side determined possibly to be weathered out
Campbell Creek Tuff with a near-vertical flow attitude (enclosed in border)....................37
Figure 13: Outcrop (A) and hand sample (B) picture of 10-BV-16. The outcrop o f this
glassy basaltic andesite is flanked on both sides by glassy flows of felsic ignimbrite
39
Figure 14: Sample 10-BV-29 in outcrop (A) and hand sample (B) part o f the Eastern suite
of FCM. (A) Shows the outcrop as bedded and very sheared, striking easterly and dipping
10°N whereas (B) shows an average abundance of vesicles present in most samples of
this suite.................................................................................................................................. 39
Figure 15: Basaltic andesite sample 10-BV-42 shows a rounded plagioclase xenocryst
displaying a high degree o f disequilibrium with the melt...................................................45
Figure 16: Basaltic andesite sample 10-BV-42 with olivine phenocrysts and iddingsite
weathering in its fractures varying in size from 1-2 mm.....................................................45
Figure 17: Highly altered andesite sample 10-BV-05 showing chloritization (light green
in ppl) on the left and seritization of plagioclase phenocrysts in xpl on the rig h t............ 46
Figure 18: Andesitic sample 10-BV-41 shows a hornblende opaque reaction rim
indicating some disequilibrium with the melt either due to decompression or
degassing................................................................................................................................. 46
Figure 19: Andesitic sample 10-BV-09 showing a large zoned plagioclase megacrysts
with sieve texture....................................................................................................................49
Figure 20: Andesitic sample 10-BV-31 showing abundant vesicles and sieve texture on
phenocrystic plagioclase........................................................................................................ 49
Figure 21: Dacitic sample 10-BV-34B showing seritized plagioclase phenocrysts, sieve
texture and are partly resorbed with reaction rims................................................................50
Figure 22: Dacitic sample 10-BV-46B showing spherulites as rounded structures which
indicates divitrification and plagioclase phenocrysts are in moderate disequilibrium
50
Figure 23: Rhyolitic sample 10-BV-48B showing spherulites and elongated brown-black
biotite with reaction rims in a glassy microcrystalline groundmass................................... 53
Figure 24: Total alkalies vs. silica diagram o f Le Bas et al. (1986)
55
Figure 25: Paleogene mafic to felsic samples plotted on a MgO, FeO*, and Na 2 0 +K 2 0
(wt%) ternary diagram of Irvine and Baragar (1971)..........................................................56
Figure 26: Silica variation diagrams for TiCh, A I 2 O 3 , MgO, CaO, Na2 0 , P 2 O 5 , FeO1,
MnO and Ca0 /Al2 0 3 all in Wt (%). Symbols are as in Figure 25...................................... 58
Figure 27: K2O vs. Si0 2 plot of Peccerillo and Taylor (1976). Symbols are as in Figure
25............................................................................................................................................. 59
Figure 28: Silica variation (wt%) diagrams for Ce, Sr, Sc, Co, V, Pb, Zr, Ba, Th, Cr, Ni
and Nb (all in ppm). Symbols are as in Figure 25............................................................... 62
Figure 29: Plot o f total iron as Fe2 0 3 vs. V. Symbols are as in Figure 25..........................63
Figure 30: Incompatible element diagrams for (a) all samples, (b) basaltic andesites, (c)
andesites and (d) dacites and rhyolites. All values are normalized to primitive mantle
(Sun and McDonough, 1988).................................................................................................64
Figure 31: Trace element ratio plot o f Ce/Pb vs. Zr/Nb. Symbols are as in Figure 25... .6 8
Figure 32: Trace element plot of Ce/Pb vs. Cs/Rb. Symbols are as in Figure 25.............. 6 8
Figure 33: Nd, Sr and Pb isotopic ratios in the FCM and surrounding areas (a, b, d), (c)
87Sr/86Sr vs. silica and (e) 143Nd/144Nd vs. silica content.....................................................69
Figure 34: K2O vs. SiC>2 diagram o f Peccerillo and Taylor (1976). FCM samples plot
higher than the Ancestral Cascades into similar and higher potassic values than the
Sulphur Springs.......................................................................................................................75
Figure 35: Ternary diagram of Th, H f and Ta discriminating between calc-alkalic basalts
(CAB), within plate alkali (WPA), within plate tholeiitic (WPT) and E-MORB, N-MORB
and island arc tholeiitic (LAT) (Wood, 1980)...................................................................... 75
Figure 36: Radiogenic isotopic 87Sr/86Sr vs. 143Nd/144Nd plot showing isotopic Sr and Nd
ranges for known studies in the Great Basin. FCM samples are shown as the yellow field,
BV = Buffalo Valley, and WGB = Western Great Basin. BV data from Wetmore (2011)
Modified from Cousens et al. (2008).................................................................................... 78
Figure 37: Primitive mantle normalized plot o f Sr/Ppmn versus 87Sr/86Sr of FCM samples
compared WGB and Sierra Nevada field areas. Block A represents Precambrian
lithopheric mantle and block B represents mantle wedge. Modified from Cousens et al.
(2008)
81
Figure 38: Shows the division o f western North America defined by the 0.706 line (large
dot-dash line). Figure from Streck et al. (1999)................................................................... 85
Figure 39: Incompatible element diagrams for the FCM (blue), Sulphur Springs (red), and
Ancestral Cascades (green). All values are normalized to primitive mantle (Sim and
McDonough, 1988)................................................................................................................ 86
Figure 40: Incompatible element patterns for FCM basaltic andesites vs. Ancestral
Cascades basaltic andesites normalized to primitive mantle (Sun and McDonough,
1989)....................................................................................................................................... 8 6
Figure 41: Ternary diagram of FeOT, MgO and AI2O 3 discriminating between different
geotectonic settings. All samples plot within an orogenic setting (island arc and active
continental margin). Fields from Pearce et al. (1977)......................................................... 8 8
Figure 42: Tectonic discrimination diagrams for rocks from the FCM compared to the
volcanic suits from East Sulphur Spring, other Eocene volcanic suites from Bingham,
Utah, and along the Carlin trend, Nevada. (A) Silicic rocks (Pearce et al., 1984). (B)
Mafic rocks (Muller and Groves, 2000). Compositions of East Sulphur Springs samples,
Bingham, Carlin, Tuscarora and Emigrant Pass are from Ryskamp et al. (2008) and
sources within.........................................................................................................................89
AAA
A A i
Figure 43: Pb isotope ratios in major terrestrial reservoirs plots for A) Pb/ Pb vs.
206pb/2°4pb m d B) 207pb/204pb yg 206pb/204pb pCM symbols same as in Figure 25. Figure
modified from White (1997)..................................................................................................91
Figure 44: REE diagrams for dacite and rhyolite FCM samples. All values are normalize
to primitive mantle (Sun and McDonough, 1988)............................................................... 97
Figure 45: Initial eNd and eSr values of granitic rocks in the northern Great Basin divided
by the groupings eugeocline, miogeocline and craton. Figure modified from DePaolo and
Farmer (1984)........................................................................................................................ 98
Figure 46: Shands index showing mainly metaluminous compositions for FCM samples.
Fields from Shand (1943)......................................................................................................98
Figure 47: REE patterns for Andesites normalized to Sun and McDonough (1989)...... 100
Figure 48: Variation diagrams comparing compositions o f FCM samples to those of East
Sulphur Spring, Carlin, and Bingham volcanic suites. Blue lines show the results o f
mixing basaltic andesite with rhyolite, and red arrows are schematic paths for fractional
crystallization of basaltic andesite. (A) Cr versus Si02. (B) Ni versus Si02. (C) Ba
versus Si02. Figures modified from Ryskamp et al. (2008) and sources within............ 103
Figure 49: Initial Sr isotope ratios vs. Rb/Sr plot for FCM samples plotted along with
Sierran Granites, Ancestral Cascades Tahoe-Reno area and Modem South Cascades.
FCM symbols are the same as in Figure 25. Sierran Granites, Tahoe-Reno and South
Cascades geochemistry from Cousens et al. (2008)...........................................................105
Figure 50: SiC>2 vs. radiogenic 87Sr/86Sr plot. L mantle = Lithospheric mantle. Symbols
are as in Figure
25............................................................................................................................................ 107
Figure 51: Plate tectonic reconstruction of the western United States model for the FCM
modified from Ryskamp et al. (2008)..................................................................................112
x
LIST OF SYMBOLS, NOMENCLATURE OR ABBREVIATIONS
FCM = Fish Creek Mountains
CAB = Calc-alkaline basalts
OIB = Ocean Island Basalt
MORB = Mid-Ocean Ridge Basalt
LREE = Light Rare Earth Elements
MREE = Middle Rare Earth Elements
HREE = Heavy Rare Earth Elements
HFSE = High Field Strength Elements (Zr, Hf, Nb, Ta)
LILE = Large Ion Lithophile Elements (K, Rb, Cs, Ba)
Chapter 1: Introduction
1.1 Introduction:
The Great Basin of Western United States is a region of Cenozoic lithospheric
extension and volcanism that includes the state of Nevada and parts of southeastern
California and western Utah. The goal o f this thesis is to investigate Paleogene volcanism
in the Great Basin focusing on the Fish Creek Mountains (FCM) and surrounding area.
The FCM, located in north-central Nevada, is a site of multiple igneous events ranging
from 35 Ma to 1 Ma, encompassing most of the igneous history of the Great Basin. The
mountains rise up to 6,512 feet above sea level and lie south o f the Battle Mountain
mining district, Nevada (Figure 1). The area of study is primarily composed o f late
Paleogene volcaniclastic and volcanic rocks that overly Paleozoic metasedimentary
rocks, that also cover a large portion of the Sierra Nevada of northeastern California as
well as western Nevada (Stewart and Carlson, 1976). These Paleogene intermediate to
felsic lava flows are related to the westward sweep of volcanism through Nevada due to
rollback of subduction of the ancient oceanic Farallon plate under North America. The
late Paleogene mafic to felsic volcanic rocks that are found throughout the FCM have yet
to be thoroughly examined to determine their origin. The suite of Paleogene volcanic
samples were collected in the summer of 2 0 1 0 along with felsic tuffacious rocks and
younger mafic volcanics as part of a larger transect across central Nevada to eastern
California. Throughout this thesis I will present the results of field work, petrographic
study, geochemical and isotopic analyses o f the volcanic rocks and describe a tectonic
model and petrologic processes responsible for their origin.
1
128®
120*
112°
T
48® - 4.
SM
Arc CM
Figure 1: Location of Fish Creek Mountains study area (enclosed in square) and
surrounding volcanic areas of the western United States (John, 2001). NNR = Northern
Nevada Rift, CRB = Columbia River basalts, SM = Steens Mountains, M = McDermitt
caldera. Dark grey area = Great Basin.
2
1.2 Justification or purpose for this study
Few detailed and modern studie has been done on the volcanology and
geochemistry o f late Paleogene (33-34 Ma) rhyolites to basalts in north-central Nevada.
Ongoing chemical analyses are being done on younger basalts (16-10 Ma and 4-1 Ma)
(Cousens) and the 24.9 Ma FCM rhyolitic tuff from the Fish Creek Mountains (Varve,
2013). An undergraduate thesis examined younger 4-1 Ma Pliocene to Quaternary flows
and cinder cones exposed in the northwestern margin of the FCM (Wetmore, 2011). By
studying these rocks, I will be able to document how much petrological and chemical
variation exists between the samples collected and determine whether or not they are
likely to be o f similar age and volcanic origin. By doing so, I will have filled a portion of
the gap in our understanding of the volcanic history and tectonic evolution o f the Western
United States.
The purposes o f this study are to:
1) Focus primarily on Paleogene felsic rhyolites to mafic basaltic andesites to
determine how and when they may have erupted within the Fish Creek
Mountains. This combined with other projects from the area will help determine
the volcanic history from California to Nevada in order to understand whether
magmatic activity in this area is linked to slab rollback and extensional tectonics.
2) To determine the mineralogical, textural, geochemical and isotopic characteristics
of the felsic to mafic Paleogene volcanic rocks within the Fish Creek Mountains
and to determine what petrologic processes are responsible for their origin.
3
3) To determine the degree of crustal contamination in the more evolved felsic
rocks.
4) To document Paleogene mantle sources and petrogenetic processes to reconstruct
evolution of magma sources through time.
5) To determine the tectonic model which reflects the interplay o f lithospheric
extension and magma generation in the mantle (asthenosphere and lithosphere)
and the crust.
1.3 Geology and tectonic setting:
1.3.1 Brief geochronology of the Great Basin
The Great Basin is a zone of Cenozoic lithospheric extension and volcanism that
includes Nevada and parts of southeastern California and western Utah forming the
widest segment of Basin and Range topography (Figure 2; Dickinson, 2006). The
thinning of the lithosphere is a result of extension which created normal faults, expressed
at the surface by sub-parallel horst and graben mountain ranges separated by down
dropped blocks which form the valleys and basins (i.e. Basin and Range). These regional
fault-bounded structural blocks are associated with extensive Paleogene volcanism and
mineral deposits. The detachment o f structural blocks involving pull apart and/or rotation
during extension may have formed widespread “plumbing systems” and allowed for
magmas and hydrothermal fluids to migrate through south-central Nevada (Kepper,
1989).
4
120*W
110*W
SO*N-
COLORAOO'
i2o*w
Figure 2: Position of the Great Basin in the Western Cordillera boxed in by the eastern
Colorado Plateau, western Sierra Nevada (SN), northern Snake River Plane (SRP) and
the southern Garlock fault (Gf) and the Mojave block (From Dickinson, 2006). CRP =
Columbia River Plateau, BM = Blue Mountains and KM = Klamath Mountains, Great
Basin Segment = Basin and Range topography. Black star represents study location.
5
The Great Basin preserves a diverse history ranging from Archean to recent. The
oldest rocks include the mantle derived Archean Wyoming Province and the
Paleoproterozoic Mojave Province which form part o f the craton. Neoproterozoic rifting
o f the craton created basement faults that were reactivated to accommodate subsequent
tectonism and act as conduits for later igneous intrusive and/or hydrothermal fluids (Cline
et al., 2005). Rifting culminated in development of a west-facing passive margin at the
edge o f the Precambrian continental crust characterized by westward-thickening
sequence of pre-mid Late Devonian sediments (John, 2001). In early Paleozoic, the
overall the subsidence rate exceeded that of sedimentation and the margin o f shallow
water marine environments shifted progressively continent-ward (Stevens et al., 1997).
Late Devonian to early Mississippian emplacement of the Roberts Mountains allochthon
over the miogeoclinal sediments is the earliest Phanerozoic contractional event to have
affected east-central California (Antler orogeny). Late Mississippian to Early
Pennsylvanian faulting created a series o f fault bounded uplifts throughout the Inyo
Mountains and other areas of east-central California. The Late Mississippian faulting also
led to the development o f a NW-trending transform fault zone that truncated the
continental margin (Dickinson, 2006). Permian to earliest Triassic contraction
deformation in east-central California occurred as a result o f convergent plate motion that
previously truncated the continental margin (Stevens et al., 1997). During the early
Triassic, sedimentation occurred during a time of relative tectonic dormancy. Middle
Triassic contraction (Sonoma orogeny) o f deformed oceanic facies was thrusted over the
Antler orogeny termed the Golconda thrust (Stevens et al., 1997). By early Late Triassic
time, east-central California east-dipping subduction zone was established beneath the
6
continental margin in east-central California producing arc plutonism, volcanism and
east-vergent contractional deformation, including the emplacement of the Sierra Nevada
Batholith and the development o f the East Sierran thrust system. This heightened period
of magmatic activity may have been caused by increased seafloor spreading rates that
produced a hotter, less dense oceanic plate that slid under the North American plate at an
increasingly lower angle. Alternatively, the rate at which the upper plate overrides the
lower plate could have also resulted in flat-slab subduction where the upper continental
lithosphere in the Western United States was moving faster than the lower oceanic
Farallon slab resulting in flat-slab subduction (Jarrard, 1986). As the younger Farallon
Plate decreased its angle o f decent, it came in contact with the underside of the
continental lithosphere where heat and pressure o f the underlying plate are thought to
have created an easterly shift of volcanism and compressional stresses continuing
throughout the rest o f the Mesozoic and therefore thickening of the continent producing
the Laramide orogeny (Figure 3A; Fiero, 1986). Magmatic front migrated east into
Colorado and did not return to Nevada until Mid-Cenozoic time (John, 2001). As the
velocity of the previously sub-horizontal subducting Farallon slab decreased, it allowed
for more time to cool, contract and sink as a result of its higher density into the hot
asthenosphere. This slab rollback is thought to have shifted volcanism in the eastern
Colorado plateau in a southwesterly direction back towards the western margin o f North
America (Figure 3B; Fiero, 1986). The arrival o f the Pacific plate replaced compression
and subduction with shear stress and a transform plate boundary which will later become
the San Andreas Fault where as the Farallon Plate continues to sink into the
asthenosphere.
7
O .W .V 5 W -
Figure 3: East-dipping continental margin subduction zone with associated periods of
heightened magmatic activity and deformation related to the subduction of the Farallon
plate under North America. A) Latest Mesozoic-earliest Cenozoic; top=80-70 Ma,
middle=75-65 Ma, bottom=60-50 Ma. B) Middle to late Cenozoic; top=45-25 Ma,
middle-20-15 Ma, bottom=5-0 Ma (Fiero, 1986; Zhang et al., 2009).
8
Cenozoic crustal extension and associated normal faulting formed first in northeastern
Nevada and northwestern Utah as previous compressional stresses of the Mesozoic and
Early Cenozoic are removed allowing the continent to spread laterally. The beginning of
extensional faulting at any given latitude in the Basin and Range generally coincided with
or immediately postdated voluminous eruptions of intermediate to silicic volcanic rocks
(Best and Christiansen, 1991). The actively extended lower crust and mantle lithosphere,
combined with crustal weakening by mantle upwelling in the wake o f the subducted
Farallon Plate, is thought to have produced widespread Basin and Range extension.
Bartley et al (1988) found crustal extension to be episodic during four periods of time;
prevolcanic before 32 Ma, early syn-volcanic (30-27 Ma), immediately post-volcanic
(16-14 Ma), and post-post volcanic Pliocene to Quaternary. Typically mantle-derived
basaltic volcanism was associated with faulting during periods of slow extension whereas
volcanism ceased during periods of rapid extension. Such mantle upwelling also has a
number of affects on the Basin and Range including: bimodal volcanism, decompression
melting, thermal expansion and softening o f the continental crust (John, 2001).
1.3.2 Great Basin igneous assemblages and volcanism
Great Basin Cenozoic igneous rocks can be divided into three assemblages: 1)
interior andesite-rhyolite (Eocene to early Miocene), 2) western andesite (early Miocene
to early Pliocene) and 3) bimodal basalt-rhyolite (middle Miocene to Holocene) (Figure
4; John, 2001).
1)
The interior andesite-rhyolite assemblage mainly contains dacite to rhyolite ash
flow tuffs and flow dome complexes with small volumes of andesitic and dacitic lava
9
flows. They are generally calc-alkaline; however much more potassic and silicic than
typical arc-related magmas (John, 2001).
Eocene magmatic activity began at around 43 Ma with the eruption o f silicic ashflow tuffs and intermediate lava flows in northeast Nevada representative o f the interior
andesite-rhyolite assemblage (Figure 4A; John, 2001). By the Oligocene, magmatic
activity of the assemblage migrated southwestward (27 Ma) from northeastern Nevada.
The youngest rocks of this assemblage erupted at about 19 Ma. Caldera complexes are
common largely in the Oligocene to early Miocene, and more than 50 calderas have been
identified in central Nevada and western Utah. Plutonic rocks, primarily granodiorite, are
exposed locally and include porphyry copper and skam related intrusions in the Battle
Mountain mining district, Nevada (John, 2001).
Heating during this period is related to the subduction of the spreading centre
between the Farallon and Kula oceanic plates that produces a slab window allowing the
upwelling of the athenospheric mantle to impinge on the base of the lithosphere (Cline et
al., 2005). Progressive rollback and separation of the Farallon plate from the base of the
lithosphere was responsible for both the onset of extension and high potassium calcalkaline magmatism in northern Nevada during the Late Eocene and eventually southcentral Nevada during the Oligocene to Miocene that lead to the formation o f basaltic
melts partly mixed with lower crustal melts (Farmer et al., 2002). This assemblage is
most similar to the mafic to felsic Paleogene samples collected in the FCM.
2)
The western andesite assemblage is a highly potassic calc-alkaline assemblage
composed mostly of intermediate lava flows, breccia and hypabyssal intrusions formed
10
along the northwestern edge of the Great Basin in western Nevada and eastern California
(Figure 4B; John, 2001). Small rhyolite intrusions and volumetrically minor basalt are
also widely distributed. This assemblage was part of the continental margin arc that was
active in and west of the modem Cascade Range. It formed in response to the subduction
o f Farallon oceanic crust beneath the continental margin of North America during the
mid- to late Miocene. Rocks of the Western andesite assemblage range in age from 22-4
Ma, the youngest of which are along the northwestern edge of the Great Basin (John,
2001).
3)
The bimodal basalt rhyolite assemblage is the youngest rock assemblage in the
Great Basin which include olivine basalt, pyroxene andesite and basaltic andesite, as well
as both sub-alkaline and peralkaline rhyolite located east o f the western andesite
assemblage (Figure 4C; John, 2001). Volcanism began in the northern Great Basin at
around 16.5 Ma and continues locally to the present day. Bimodal volcanism in the northcentral Great Basin may have initially been concentrated at the north end o f the Northern
Nevada Rift. Volcanism from this rift lasted from 16.5 Ma to 15 Ma and extends 500km
from the Oregon-Nevada border to south-central Nevada. Generally, intermediate rocks
(siliceous andesite, dacite and trachydacite) are common in the central and northern parts
of the Northern Nevada Rift but are generally uncommon everywhere else. The older
phases of this assemblage formed in a back-arc environment related either to back-arc
extension or to the encroachment of the Yellowstone mantle plume on the crust at 16.5
Ma along the Nevada-Oregon border. The younger phases formed during continental
extension unrelated to subduction, perhaps as a result of lithospheric extension over a
mantle plume (John, 2001).
11
Figure 4: Maps showing the general
distribution of Cenozoic volcanic
assemblages outlined In dark grey.
A) Interior andesite-rhyolite assemblage
showing 27 Ma timeline illustrating a
southwestward sweep of magmatism.
B) Western andesite assemblage.
C) Bimodal basalt-rhyoiite assemblage.
Black stars represent study location.
(Figure modified from John, 2001)
12
1.3.3 Great Basin Tectonics
The subduction of the oceanic Farallon plate and associated rift (the East Pacific
Rise) played an important role in the formation of the Great Basin. Tectonics in the
western United States can be broken up into two main time periods; (1) latest Mesozoic
to earliest Cenozoic and (2) middle-late Cenozoic.
1)
During the latest Mesozoic to earliest Cenozoic the creation o f the subduction
zone that consumed the oceanic Farallon plate lead to an easterly migration o f arc
volcanism to the Colorado Plateau.
By early Late Triassic time, east-central California was influenced by an eastdipping continental margin subduction zone. As a result, the Mesozoic was a period of
heightened magmatic activity as the arc evolved from Late Triassic to Late Cretaceous
time. Abundant Triassic (220-210 Ma) and larger Jurassic (152 Ma) plutons and mostly
Cretaceous (148 Ma) major dyke swarms intruded the region, where as volcanic
complexes (222-98 Ma) accumulated on the surface (Stevens et al., 1997). However,
Sierra Nevada arc activity ended as a significant increase in spreading rates produced a
hotter and therefore less dense oceanic plate that slid under the North American plate at
an increasingly lower angle producing an easterly wave o f volcanism until flat-slab
subduction. Flat-slab subduction compressed the lithosphere and created the Laramide
orogeny (Figure 3A; Fiero, 1986). Magmatic activity migrated east into Colorado and did
not return to Nevada until Mid-Cenozoic time, possibly due to crustal thickening during
the Late Cretaceous Sevier and Laramide orogenies associated with low-angle
13
subduction, accretion of island-arc terranes and progressive contraction o f the
miogeocline from west to east.
2)
Middle to Late Cenozoic crustal extension dominated the tectonic history of the
Great Basin during the development of several styles of faulting. Beginning in the late
Eocene, stress relaxation and mild extension, characterized by multiple sets o f normal
faults and detachment faults in more deeply exposed terranes, affected much of the Great
Basin whereas more extreme extension occurred in the Miocene. Generally, areas of
extension formed first in northeastern Nevada and northwestern Utah. This rapid
extensional period was followed by an increased angle o f the subduction of the oceanic
plate leading to a shift of volcanism from easterly back to a westerly direction. As the
impact velocity o f the previously sub-horizontal subducting Farallon slab decreased, it
allowed for more time to cool, and sank as a result o f higher density into the hot
asthenosphere (slab rollback). This triggered melting of the overlying lithosphere due to
previously released volatiles and caused magmatism at the surface (Best and
Christiansen, 1991). As the plate continued to steepen this locus of volcanism would shift
westerly to its most recent expression as late Paleogene volcanism in the Sierra Nevada in
eastern California (Cousens et al., 2008). As the plate sank deeper it eventually broke off
causing the upwelling of the asthenosphere (Figure 3B; Fiero, 1986)
At about 16 Ma, more widespread Basin and Range extension began, producing
alternating basins and ranges spaced 20-50 km apart (John, 2001). The extension that
created the Northern Nevada rift and continued to form the present day Basin and Range
topography probably began in the mid-Miocene (Ressel and Henry, 2006). Local, rapid
large magnitude extension continued after the initiation of Basin and Range faulting.
14
Recent estimates for the total amount of Cenozoic extension in the Great Basin are
between 100 to 250 percent. More extensive Basin and Range extension occurred in the
mid-Miocene in contrast to the Eocene when the North American continent came into
greater contact with the Pacific plate creating the San Andreas Fault in between remnants
o f the subducted Farallon plate termed the Gorda and Cocos plates. This is thought to
have removed compression between the Farallon plate and the continent, giving the
continent room to spread (John, 2001).
1.4 Local geology (Nevada)
Paleogene igneous activity in Nevada has been suggested to be the result of
magmatism due to the subduction o f the Farallon Plate and associated spreading centre
under North America (Best and Christiansen, 1991). Volcanism can be broken up into
three periods on the basis o f K-Ar dating (McKee and Silberman, 1970). At about 50 Ma,
a magmatic front began migrating southwestward across southern Idaho, central Oregon
and into northern Nevada and Utah. Intermediate, arc like basaltic andesites through
dacites dominated volcanic activity in northeastern Nevada in between 45 to 36 Ma.
From 34 to 24 Ma, the extrusions o f ash flow sheets covered large parts o f southern and
eastern Nevada, termed the ignimbrite flare up (Fiero, 1986). After about 24 Ma
magmatism changed gradually in composition and style of eruption. Mafic cinder cones,
tuff cones, low shield volcanoes, isolated lava flows, and viscous rhyolitic domes became
more characteristic and smaller calderas and ash-flow sheets became less common
(Christiansen, 2010). The final period of Paleogene magmatism lasted from 16 to 10 Ma,
including basalt and basaltic andesite flows along with intrusive rhyolite flow dome
complexes found around the FCM and are related to the volcanism of the Snake River
15
plain province to the north (McKee and Silberman, 1970). After approximately 10 Ma,
the magmatic rocks became increasingly bimodal, with the appearance o f basalt and the
disappearing of intermediate magmas. Volcanism then migrated westwards towards into
Sierra Nevada as a result o f progressive sinking and southwestward rollback o f a
shallowly dipping subducting slab. This resulted in widespread dehydration o f the
subducted lithosphere and generated voluminous mantle-derived magma which intruded
and differentiated in the crust (Christiansen, 2010). The inflow of asthenospheric mantle
beneath the Great Basin combined with the development o f a transform boundary and
regional extension resulted in decompression melting of asthenospheric and lithospheric
mantle. A portion of this magma stagnated in the lower crust, then re-melted and
differentiated to create subalkaline and peralkaline magmas o f anorogenic affinity in
contrast to the predominantly calc-alkaline affinity o f the central Great Basin. These
mafic lavas are typically alkali basalts (lack negative Nb-anomalies) and for the most
part, initial 87Sr/86Sr isotopic ratios are lower where as ,43Nd/,44Nd isotopic ratios are
higher than the older arc magmatism (Christiansen, 2010).
Tuffaceous volcanics in central Nevada can be distinguished in three different
units; the upper Bates Mountain tuff (23.1 +/-1.7 Ma), the Caetano tuff (31.2-33.3 +/- 2.7
Ma) and the FCM tuff (24.9 Ma). The upper Bates Mountain tuff is concentrated
northwestward with only a few remote outcrops recognized in the northern part of the
FCM. The other two are welded tuffs that make up the Paleogene igneous section in
north-central Nevada. The Caetano tuff is concentrated in a west-trending belt from the
northernmost Toiyabe Range to the west edge of the central part of the Shoshone Range
but also crops out at the northern end of the FCM. The youngest FCM tuff is exposed
16
throughout the FCM but the outflow sheet is limited to a few miles away from the FCM
(McKee and Silberman, 1970).
Mesozoic and Paleogene intrusive suites are much less extensive and most
common in the northern part of Nevada (McKee and Silberman, 1970). Based on K-Ar
dating, the intrusive rocks range in age from Oligocene to Cretaceous (McKee and
Silberman, 1970). Most of the intrusive suites are composed of fine-grained porphyritic
quartz diorite to quartz monzonite. Granite Mountain (49-30 Ma) is the largest Paleogene
intrusive body in the study area being more than twice to size of any others nearby. It is
composed o f medium to coarse grained hypidiomorphic granular rock similar to the large
Mesozoic plutons of central Nevada (McKee and Silberman, 1970).
1.5 Components in continental subduction zones
The subduction of the oceanic Farallon plate and associated mid-ocean rift played
an important role in the formation o f the Great Basin. Paleogene Nevada igneous activity
has been suggested to be the result of arc volcanism (Best and Christiansen, 1991). There
are four main components in subduction zones; role of the subducting slab, the subducted
sediments, mantle wedge and fluids overlying the subducting oceanic crust, and the
overriding plate (mantle and crust).
1) Role of the subducting slab
The asthenosphere is 1-2% less dense than the lithosphere, and is what drives
plate motion and subduction. As the lithosphere ages, the density increases and the
subducting slab thickens and cools plunging beneath the continent at a higher angle. The
differences between lithospheric age and dip of the subducting slab are reflected in
17
different subduction styles flanking the Pacific plate. Alternatively to the Pacific Rim, the
eastern younger, warmer and less dense plate relates to a shallow angle and often failed
subduction zones (Davies, 1999). Fluids driven from the subducting slab are added to the
mantle wedge before they give rise to primary magmas.
2) Role of sediment in subduction
Variations among mobile LILE trace elements (K, Sr, Ba) and more specifically
Be 10 have been interpreted to indicate the presence o f sediment addition to the mantle
wedge (Plank and Langmuir, 1993). The limited range in compositions present in
lithopheric mantle and oceanic crust differs with the compositional variation o f the
sedimentary packages being subducted and often provides a unique geochemical input
into the zone of subduction. For instance, Peru-Chile subduct carbonates where as the
Aleutians subduct terrigeneous sediment (Rea and Ruff, 1996). Lead isotopes can also
provide a possible insight into the role of sediment in magma geochemistry within the
mantle wedge and may show steep arrays on the 208 Pb/204Pb and 207Pb/204Pb vs.
206Pb/204Pb
plots between MORB and/or overlapping with local sediments from the
subducting plate (Hamilton, 1994).
3) Mantle wedge and fluids:
The mantle wedge is the part of the asthenosphere that is located directly above
the subducting slab and below the overriding plate. This is where components from the
descending slab are commonly mixed with the converting mantle to generate magmas
and eventually produce new continental crust. Melts generated in the mantle wedge are
produced due to in flux of slab-derived fluids and lowering o f the solidus of the mantle,
18
as oppose to decompression melting associated with mantle plumes and mid-ocean ridges
(Brenan et al., 1995). Aqueous and carbonic fluids are continually released from the
subducting slab, from the crust, subducted sediments, and by mineral dehydration leading
to metasomatization o f the overlying mantle wedge. Some elements, including the large
ion lithophile elements (LILE; K, Rb, Cs, Sr, Ba, Pb, and U) are transported in hydrous
fluids, where as high field strength elements (HFSE; Y, Zr, Hf, Nb, Ta), the rare earth
elements (REE), Ti and the transition elements remain relatively immobile (Brenan et al.,
1995). Distinct enrichments of LILE relative to HFSE are observed on trace element plots
of arc lavas, which reflect the mixing of LILE carried in the aqueous fluid with the
mantle wedge supply of HFSE.
4) Role of the lithospheric mantle in overriding plate
The composition o f the continental mantle lithosphere varies with age (i.e.
density, thickness and fertility). Archean mantle lithosphere contains relatively low FeO
abundances linked to komatiite extraction making it less dense. Typically, post-Archean
mantle lithosphere is more enriched in FeO, and hence dense enough to be delaminated
and incorporated into asthenospheric source regions of oceanic basalts and is thought to
be similar to spinel peridotite inclusions in alkali basalts. The above process is therefore
more likely to produce continental flood basalts such as the Columbia River Flood basalts
(Hawkesworth et al., 1990). Basalts erupted at active continental margins have Nb, Ta, Zr
and H f components that are not seen in oceanic island arcs and assumed to be from
enriched metasomatized subcontinental lithosphere (Pearce, 1983). The subduction
process will generally provide Rb, K, Ba, Th and Sr where as the lithosphere contributes
portions of LREE and P and all o f its Ta, Nb, Zr, Hf, Ti, and HREE (Pearce, 1983).
19
During flat-slab subduction of the Farallon plate under North America (70-45
Ma), water may have been released from the down-going slab resulting in the hydration
o f the North American mantle lithosphere. The cooling as a result of the displacement of
asthenospheric in flow and the passage o f cold oceanic lithosphere beneath the trapped
overriding continental lithosphere resulted in a period of ceased magmatism in Western
U.S because the mantle material above the flat slab was not hot enough to form melt
(Humphreys et al., 2003). As flat-slab subduction ceased, the removal o f the Farallon slab
would have exposed the hydrated mantle lithosphere to the underlying hot asthenosphere,
possibly melting this upper mantle lithosphere and resulting in the rhyolitic ignimbrite
flare-up (Humphreys et al., 2003).
1.6 Degree of crustal contamination
The average crustal compositions above Mesozoic and Cenozoic subduction
zones are primarily intermediate (andesitic) with high Rb/Sr ratios relative to Bulk Earth
(Ellam & Hawkesworth, 1988). As mantle-derived basaltic melts enter the crust, they
undergo fractional crystallization or solidify in deeper regions of the crust until they are
remobilized (Ellam & Hawkesworth, 1988). Therefore the mantle to crust flux is
differentiated into more intermediate to felsic material that rises into the upper crust
where as more mafic cumulates and/or remains at lower crustal depths. Andesites are
therefore more likely to have been derived from fractional crystallization and crustal
contamination processes rather than as primary melts of the mantle (Ellam &
Hawkesworth, 1988).
20
In continental subduction zones, continental crust up to 70 km thick may be
present. Consequently the geochemistry o f the Paleogene mafic magmas is expected to
have been modified as they ascended to the surface, especially through Proterozoic and
Paleozoic sedimentary rocks. The crust is generally composed of low density felsic rocks
enriched in incompatible elements. As the denser yet thermally buoyant mantle derived
mafic magma works its way up through the crust it often undergoes differentiation in
areas of magma stagnation. Denser minerals precipitate out o f the melt and sink, the
residual melt becomes more siliceous and heat latent of crystallization is released into the
country rock. This, combined with the low melting temperature of the crust, results in
partial melting of the cmst which incorporates or contaminates the magma, altering the
geochemical and isotopic components towards the crustal composition (Ida, 1983). The
felsic minerals of the continental crust can include large amounts of incompatible
elements that will preferentially partition out of the cmst and into the melt (Rea and Ruff,
1996).
21
Chapter 2: Methodology
2.1 Field Methods
From July 12 - July 18, 2010, samples were collected from the Fish Creek
Mountains field area, with the guidance o f the geologic map o f North-Central Nevada
(Stewart and Carlson, 1976), geologic map of McCoy Mining District (Emmons and Eng,
1995), geology of Golconda Canyon of the Southern Tobin Range (Gonsior, 2006) and
Nevada state maps. Samples were collected from outcrops that generally represented a
lava flow; however, where flows were poorly exposed, samples were collected from
voluminous float determined to be part of the flow. For the most part, all samples were
weathered to a brown-red surface: float samples tend to be more weathered. Weathered
surfaces were removed for thin section and geochemical analysis by rock hammer in the
field. GPS locations were recorded using a Magellan handheld unit. A map o f the area
with sample locations is presented in Figure 5 and rock descriptions with sample number
and locations are presented in Appendix A.
2.2 Petrography
All but one of the 29 samples collected in the Fish Creek Mountains were cut into
thin sections and powdered for chemical analysis. H I0-53 was the only sample not made
into thin section. It was collected by Chris Henry at the Nevada Bureau o f Mines and
Geology to be dated by Ar-Ar methods, and only the crushed powder was available. To
prepare thin sections, each sample was loaded into the diamond rock saw and crosssectionally cut into 1 cm thick disks. Three disks were cut per sample, one for a thin
section and the other two were to be crushed for geochemical analysis. The disc cut for
22
thin section preparation was made into a 3x5 cm puck using a smaller hand rock saw. To
obtain a smooth surface, the pucks were then polished on a spinning rock polisher using a
series of fine grit sands. A total 28 thin sections were produced for microscopy. Thin
sections were studied using 4-1 Ox magnification both in plain and cross polarized light on
a Mel Sobel light microscope. A petrographic summary is in Appendix B.
2.3 Geochemistry powder preparation
The remaining two disks that were prepared on the diamond rock saw were then
trimmed on the smaller wet saw to remove the weathered edges. The discs were then
broken into smaller fragments using a hammer in order to fit them into the steel plated
Braun Chipmunk crusher. The crusher yielded chips ranging in size from 1 mm to 2 cm
which were pulverized in a Rocklabs steel ring mill for approximately 45 seconds or until
they were an ultrafine powder. The steel head was cleaned for each sample using a run o f
silica sand to prevent cross-contamination. Soft steel milling and crushing may have
added small amounts of iron to the rocks, however the expected contamination o f 0.2-0.3
wt % is considered to be insignificant with respect to this study (Iwansson and
Landstrom, 2000). Minor amounts of Cr were added to each sample from the steel head,
but tests of samples crushed in agate versus steel showed that no other elements were
added during crushing. Two vials o f powder were produced; one sent for ICP-MS (trace
element) at the Ontario Geological Survey, and the other used for XRF analysis of major
element oxides and selected trace elements at the University o f Ottawa as well as for
radiogenic isotope geochemistry at Carleton University.
23
2.4 Major element geochemistry - X-Ray Fluorescence (XRF)
The analyses of major and trace elements in geological materials by X-ray
fluorescence (XRF) is made possible by the behavior of atoms when they interact with xradiation. An XRF spectrometer works for the reason that a sample is illuminated by an
intense incident x-ray beam, causing excitation of atoms in the sample and release of xray spectra that varies depending on the composition of the sample (La Tour, 1989).
Samples were analyzed at the University o f Ottawa X-Ray Fluorescence facility
where the concentrations o f Si0 2 , TiOi, Fe203‘, MnO, MgO, CaO, Na 2 0 , K 2O and P2O 5
were determined. In addition, Ba, Co, Ga, La, Ni, Pb, Rb, Sr, Th, U, V, Y, Zr, Nb, Cr, Ce
and Nd were also determined. The XRF at the University o f Ottawa uses a fused disc
technique. Loss on ignition (LOI) was calculated from the weight difference before and
after fusion o f sample and flux at 1000°C. The 28 samples were run with the internal
standard 00-LT-2 and average deviation to ensure precise results. Data and standards are
listed in the Appendix C.
2.5 Trace element geochemistry - Inductively Coupled Plasma Mass Spectrometry (ICPMS)
Samples were sent to the Ontario Geological Survey’s Geo Labs in Sudbury for
acid-dissolution ICP-MS analysis where the 29 samples were run along with internal
standard of (00-LT-2) and average deviation to ensure precise results. Concentrations o f
Ba, Be, Bi, Cd, Ce, Co, Cr, Cs, Cu, Dy, Er, Eu, Ga, Gd, Hf, Ho, La, Li, Lu, Mo, Nb, Nd,
Ni, Pb, Pr, Rb, Sb, Sc, Sm, Sn, Sr, Ta, Tb, Th, Ti, TI, Tm, U, V, W, Y, Yb, Zn and Zr
24
were determined by ICP-MS. Geochemical results are presented in appendix D and data
plotted in GCDKit 3.0 (Janousek et al., 2006).
In general, the samples used for ICP-MS are introduced into an argon plasma as
aerosol droplets where the plasma dries the aerosol, dissociates the molecules, and then
removes an electron from the components forming singly charged ions. These charged
ions are then directed into a mass filtering device known as the mass spectrometer. Only
one mass to charge ratio will be allowed to pass through the mass spectrometer from
entrance to exit at any given time. When exiting the mass spectrometer, ions strike the
first dynode of an electron multiplier serving as a detector. The impact o f the ions
releases a surge of electrons that are amplified until they become a measurable pulse. The
software compares the intensities o f the measured pulses to those from known standards,
which make up the calibration curve, in order to determine the concentration of the
element.
It is typically only necessary to measure one isotope for each element measured
since the ratio o f the isotopes, or natural abundance is fixed in nature. On the other hand,
naturally occurring lead originated from two sources; some was placed here when the
Earth was formed and some is the result of the decay of radioactive materials. Therefore,
lead isotope ratios may vary depending on the source of the lead. In order to accurately
measure the concentration of lead in a sample, it is necessary to sum several o f the
isotopes available.
25
2.6 Isotope Geochemistry —Thermal Ionization Mass Spectrometry (TIMS)
TIMS was conducted at the Isotope Geochemistry and Geochronology Research
Centre (IGGRC) at Carleton University. Strontium (Sr) and Neodymium (Nd) isotopic
geochemistry were done on 24 samples. These samples showing the least alteration based
on LOI values less than three. Lead (Pb) isotope geochemistry was then done on 12
samples chosen based on location within the FCM and unique petrographic
characteristics. CHUR values used for Nd = 0.512638.
For each sample, approximately 100 mg were weighed and put into Teflon screwcap beakers. The samples were dissolved in 50% HF-12N HNO 3 for two days before
being dried down to a moist paste. 7N HNO3 and subsequently 6 N HC1 were added, with
dry down after each step to ensure that there was no undissolved residue remaining on the
sample. Sr-Nd samples were finally dissolved in 2.5N HC1, and Pb samples in IN HBr.
Lead was separated out of the sample in polyethylene columns containing anion
resin using IN HBr flush the other elements, then 6 N HC1 to remove the lead. The
collected lead solution was then dried and dissolved in IN HBr. A second pass of this
procedure was then done to ensure purity of the collected lead.
Each Sr-Nd sample was pipetted into a column containing cation resin. Strontium
was removed using 2.5N HC1 and the rare earth elements were then removed using 6 N
HC1. The rare earth element solution was dried down and then dissolved in 0.26 HC1.
Next, the solution was pipetted into columns containing coated Teflon powder, and the
neodymium was removed using 0.26N HC1.
26
Samples were analyzed for their Pb, Sr, and Nd isotopic composition utilizing a
ThermoFisher Triton TI thermal ionization mass spectrometer at Carleton University
(techniques o f Cousens, 1996). All Pb mass spectrometer runs are corrected for
fractionation using NIST SRM981. The average ratios measured for SRM981 are
206pb/204pb = 16.890 ± 0.009, 207Pb/204Pb = i 5 .4 2 6 ± 0.009, and 208Pb/204Pb =
36.494 ± 0.031, based on 25 runs between September 2008 and May 2012. The
fractionation correction is +0.13%/amu (based on the values o f Todt et al., 1984). Sr
isotope ratios are normalized to
8 6 Sr/88Sr
= o . l 1940. Two Sr standards are run at
Carleton University, NIST SRM987 ( 8 7 Sr/86Sr = 0.710239 ± 14, n=30, Sept. 2008 May 2012) and the Eimer and Amend (E&A) SrC03 ( 8 7 Sr/86Sr = 0.708012 ± 15, n=14,
Sept. 2008-May 2012). Nd isotope ratios are normalized to 146^^/144^(1 = 0.72190.
Thirty runs of an internal Nd metal standard yield 143;Nd/144Nd = 0.511823 ~ 12 (Sept.
2008-May 2012), equivalent to a value for the La Jolla standard of0.511852. All quoted
uncertainties are 2 -sigma standard deviations of the mean.
Age-corrected isotopic data can be found in the Appendix E. The samples were
run with lab standards to check for precision. The error listed in the table along with the
samples is the internal which is equal to 2 standard deviations of the mean. Precision* is
the internal error o f the run, the variability caused by temperatures and voltages used to
excite the sample. The external error is the reproducibility o f the run, which is checked
by running a standard with each wheel run on the mass spectrometer. This standard is
added to, and checked against, the average of standards run in the past year to assure the
27
machine is running accurately. The standard for each isotope ratio is listed below the data
table in the Appendix E.
eNd values were calculated using the following formula.
eNdT = f 143Nd/ l44N<L,m„i,. —143Nd/
144NdrHi,p )
i«N d/,44NdcHUR
28
* 10,000
T
= 34 Ma.
Chapter 3: Geology and Field Relationships In the Fish Creek Mountains
3.1 Physical Observations
Paleogene flows o f the FCM and surrounding area were either overlain or
surrounded by the 24.9 Ma FCMT, 33.8 Ma Caetano tuff, 28.8 Ma Campbell Creek tuff
or younger 16-10 Ma basalt and basaltic andesite flows related to the volcanism o f the
Northern Nevada Rift to the North. Flows have also filled paleocanyons in at least two o f
the sample locations.
Photographs of representative samples are included from Figure 6-14. Appendix
A summarizes the basic geological information from the field for the 29 samples used in
this study where as the thin section petrography of the volcanic rocks is summarized in
Appendix B for all samples.
The lavas range in composition from basaltic andesites to rhyolites. Paleogene
lavas from the FCM occur mainly in 4 different geographical locations; furthest west
within the Golconda Canyon of the Tobin range, a western FCM suite, an eastern FCM
suite within the Horseshoe Basin, and a southern suite taken from the Shoshone range
immediately south o f the FCM (Figure 5). Generally, the FCM western flow trends northsouth, where as the flows o f the Tobin Range, Shoshone range, and eastern outcrops
occur in clusters intermingled with tuff and younger syn to post-extensional basalt.
Paleogene volcanic samples taken from the Tobin Range to the west of the FCM
(Figure 5) (samples 10-BV-39 through 42) are all classified as basaltic andesites (Le Bas
et al., 1986) except for sample 10-BV-41 which is an andesite.
29
,.»i---------,---------1
---------1
---------1
---------1
---------1
—
-117.75
-117.55
-117.35
-117.15
Figure 5: Rock sample locations for samples collected during the 2010 field season
plotted as blue crosses on the topographic map of the FCM and surrounding area. They
are divided into four main geographical areas (encircled) Red crosses represent 24.9 Ma
rhyolitic tuff. (Modified from Fish Creek Mountains and Edwards Creek Valley USGS
topographic maps, U.S. Geological Survey, 1:100 000 scale).
30
They were all sampled from within the Golconda Canyon area of the southern Tobin
Range which consists of a pre-Paleogene east-west trending paleo-valley in filled and
overtopped by Paleogene volcanic and sedimentary units (Gonsior, 2006). Sample 10BV-42, located at the west end of the Golconda Canyon is overlain by 33.8 Ma Caetano
tuff. This sample comprises a vertically jointed outcrop typical of a thick mafic flows that
are common in the western FCM. Other Paleogene units found in the Golconda Canyon
but not sampled include a biotite rhyolitic ignimbrite (33.28 +/- 0.22 Ma) and a thick
sequence of andesitic rocks overlying the Caetano tuff in most exposures (32 Ma)
(Gonsior, 2006). Basaltic andesite samples, excluding sample 42, were all vesicular and
exposed in weathered outcrops. In general these samples all have a glassy aphanitic
groundmass with up to 20-25% euhedral plagioclase phenociysts l-5mm and 10-15% up
to 7mm hornblende phenocrysts with minor pyroxene (Figure 6 ). Sample 41, the
hornblende andesite, taken from an outcrop that shows vertical joints, lithic clasts and
boulders on top yet has a volcanic matrix filled with crystals (Figure 7). This flow is
probably volcaniclastic (collapsed dome or debris flow). Most samples contained
weathered mafic minerals and iron oxide stained fractures. Phenocrysts range in size
from millimeters to no longer than
1
centimeter.
Lavas from the western FCM found in Jersey Canyon (Figure 5) range from
andesite through to rhyolite and can be further subdivided into vesicular and nonvesicular samples. Samples 10-BV-07 through 10-BV-l 1 are all non-vesicular, massive,
and mostly porphyritic andesites taken from columnar jointed outcrops typical o f thicker
flows (Figure 8 ). Sample 10-BV-09 strikes 294 with a dip o f 70° N and was taken in
close proximity to samples 7 to 11. The remainder o f the samples were taken from
31
different types o f vesicular flows and tended to be more evolved dacites to rhyolites.
Sample 10-BV-12A and B taken from the upper portion of Jersey Canyon are most likely
rhyolitic volcanic breccias where as sample 10-BV-34A and B were embedded in a
section of FCMT as rounded dacite lithic clasts. Sample 10-BV-06 located just south of
sample 34 is suspected to have been part of a lava flow top or a’a’ (Figure 9). Most
samples exhibited trachytic texture. Similar to the Tobin Range samples, all western
FCM samples also tend to have a glassy aphanitic groundmass with plagioclase and
hornblende microphenocrysts.
Southern samples were collected just south of FCM part of the Shoshone
Mountain range and next to the Red Butte area (Figure 5). These flows range in
composition from andesite through to rhyolite; however this region typically includes
more evolved rocks, commonly dacites through rhyolites, and can be subdivided into
samples taken from non-tilted columnar jointed outcrops versus those from tilted fissile
and platy outcrops. The andesites are more commonly found in jointed outcrops in
contrast to the dacitic rocks usually found in platy ones. Sample 10-BV-45 showed crude
columnar jointing o f fresh lava and contains trace anhedral olivine with minor plagioclase
(Figure 10). Sample 10-BV-48A, taken from a dacitic cone, contains shear zones typical
of more viscous flows and a very glassy matrix similar to the jointed andesites. The rest
of the samples were taken from platy outcrops commonly containing porphyritic
hornblende (1mm) and 15 % euhedral plagioclase up to 5mm long. Other samples contain
trace amounts of biotite, olivine and pyroxene.
32
(A)
(B)
Figure 6 : Typical basaltic andesite representative sample (10-BV-39) from the Tobin
Range in outcrop (A) and hand sample (B). The outcrop is fairly weathered and in hand
sample is vesicular with a glassy aphanitic groundmass with up to 25% plagioclase (15mm) and 10% hornblende (7mm) phenocrysts with minor pyroxene.
Figure 7: Sample 10-BV-41 in outcrop (A) and hand sample (B). This andesite was taken
from an outcrop that could have been a collapsed dome or a debris flow that shows
vertical joints, lithic clasts and boulders on top yet has a volcanic matrix filled with
crystals.
33
Figure 8 : Sample 10-BV-10 in outcrop (A) and in hand sample (B) characteristic of nonvesicular flows from the western FCM suit. They are mostly porphyritic andesites taken
from columnar jointed outcrops.
(A)
(B)
Figure 9: Sample 10-BV-06 in outcrop (A) and hand sample (B). This is the most
northern expression of Paleogene volcanics in the western FCM suite and is suspected to
have been part of a lava flow top or a’a. Characteristically it is more vesicular and
weathered compared the porphyritic andesite samples taken further south.
34
Dacitic sample 10-BV-47 was taken from a fissile, west-dipping bedded outcrop, and the
float down the hill from the outcrop has clasts of dacite mixed with glass and other clasts
and sediment that may have originally been flow top material (Figure 11). Sample 10BV-47, located on the ridge north of Red Butte includes euhedral plagioclase and biotite
in a crystalline matrix. Sample number 10-BV-48 A and B were taken from a dacitic to
rhyolitic cone which appeared to have lava tunnels or splash up the side, determined to be
28.9 Ma Campbell Creek Tuff. Here the tuff appears to flow up along the margin of the
dacite cone with a near-vertical flow attitude, indicating that the cone is older than the
Campbell Creek Tuff (Figure 12). Sample 10-BV-48B has spherulites which are
indicative of divitrification (formation o f minerals from glass). The sample contains
abundant plagioclase and biotite. In hand specimen southern Shoshone Mountain Range
samples all have a glassy aphanitic groundmass with trachytic texture. Therefore southern
flows are typically platy in outcrop and seem to be more evolved than western and
eastern flows.
Eastern flows are found on the east side of the FCM enclosed in the area known
as Horseshoe Basin (Figure 5). All rock samples are andesites except for 10-BV-16,
which is a basaltic andesite. 10-BV-16 has trachytic texture around olivine phenocrysts
and it is restricted to this part of the basin, flanked on both sides by glassy flows of
Pinnacle Mountain rhyolite flow dome complex (32.9 +/- 0.9 to 34.1 +/- 0.9) (Emmons
and Eng, 1995) (Figure 13). Since this basaltic andesitic large flow is clearly overlain by
a vitreous flow, it is most likely that the glassy unit is basal vitrophyre o f Pinnacle
Mountain rhyolite flow and the upper peak is devitrified Pinnacle Mountain tuff.
35
(A)
(B)
Figure 10: Sample 10-BV-45 in outcrop (A) shows crude columnar jointing o f fresh lava
and in hand sample contains trace anhedral olivine with minor plagioclase (B) sampled
from the Shoshone Range.
Figure 11: Sample 10-BV-47 in outcrop (A) and hand sample (B) characteristic of
samples taken from platy and fissile outcrops of the Shoshone range. These samples
commonly contain porphyritic hornblende (1mm), plagioclase up to 5 mm long and
biotite in a crystalline matrix. The fissile dacitic beds of sample 10-BV-47 dip west and
are located on the ridge north of Red Butte. Hand sample also shows flow texture of
plagioclase laths.
36
Figure 12: Dacitic-rhyolitic cone of sample 10-BV-48 from the Shoshone range which
appeared to have lava tunnels up the side determined possibly to be weathered out
Campbell Creek Tuff with a near-vertical flow attitude (enclosed in border).
In general, eastern flow rocks are typically highly vesicular (more so than any other) and
have a glassy aphanitic groundmass with xenocrysts and megacrysts of euhedral
plagioclase (25%) and hornblende (5%) up to 6 mm. Sample 10-BV-29 and 10-BV-30
overlie each other and are located at the southern end of Horseshoe Basin. Sample 10BV-29 is bedded and very sheared, striking easterly and dipping 10°N (Figure 14).
Samples 31 and 32 were also sampled in close proximity. In sample 31, there is no
observed bedding, is more vesicular (round to elongate) and less crystalline than any
other in this group, whereas sample 32 appears to be non-vesicular occurring in a jointed
outcrop similar to more viscous flows of the western and southern sample locations.
Therefore, with the exception of sample 32 and the localized basalt o f sample 16, easterly
samples are mainly hornblende andesitic flows, which are less viscous or thick due to the
abundance of vesicles in this assemblage.
To conclude, sampled flows from the Tobin Range (Golconda Canyon) are
generally vesicular basaltic andesites; however more evolved rocks also exist but were
not sampled. Flows from the western Jersey Canyon are generally thicker and possibly
more viscous due to their lower abundance o f vesicles typically found in jointed outcrops,
and more evolved dacite to rhyolite were also observed but are less extensive than the
other areas of FCM. Southern flows of the Shoshone Mountains contained the most
evolved rocks, typically dacite to rhyolite, often found as platy and fissle outcrops.
Lastly, the eastern flows within Horseshoe Basin were mostly hornblende andesites and
contained the greatest proportion of vesicles implying a less viscous flow or gas-rich
magmas.
38
Figure 13: Outcrop (A) and hand sample (B) picture o f 10-BV-16. The outcrop of this
glassy basaltic andesite is flanked on both sides by glassy flows of felsic ignimbrite. It
has flow textures around olivine phenocrysts and is restricted to this part of the
assemblage.
Figure 14: Sample 10-BV-29 in outcrop (A) and hand sample (B) part o f the Eastern suite
of FCM. (A) Shows the outcrop as bedded and very sheared, striking easterly and dipping
10°N whereas (B) shows an average abundance of vesicles present in most samples of
this suite.
39
Chapter 4: Geochronology
4.1 Introduction
A number of Ar-Ar ages have been established from labs at the IJSGS sampled
from the FCM area in addition to a few from the surrounding rock formations of the
Buffalo Valley area and are presented below.
4.2 Previous dating
Paleogene andesitic ages from the surrounding rock units of the FCM have been
determined and are as follows; hornblende andesite sample 08-DJ-126 from Mt Caetano
(35.2 Ma), andesite sample H05-52 from the Toiyabe Range (35.2 Ma), glassy
clinopyroxene-plagioclase andesite porphyry sample 09-DJ-85 from the Fye Canyon
Volcanics (35.5 Ma), and 09-DJ-87 porphyry homblende-plagioclase andesite also from
the Fye Canyon Volcanics with an age o f 35.5 Ma (John et al., 2008).
Four ages were determined in the Golconda Canyon, part of the Tobin Range by
Gonsior (2006); biotite rhyolitic ignimbrite sample TR05-26 (33.28 +/- 0.22 Ma),
ignimbrite TR-9 (33.03 +/- 0.25 Ma), rhyolitic ignimbrite TR05-21 (24.95 +/- 0.17 Ma)
and a younger basalt sequence TR-77 (14.10 +/- 0.12 Ma). Surrounding tuffaceous units
include the Caetano and Campbell Creek tuffs with ages of 33.8 Ma and 28.8 Ma
respectively. The thick sequence of andesitic rocks overlying the Caetano tuff in most
exposures in the Tobin Range has been dated by McKee et al., (1971) with a K-Ar age of
approximately 32 Ma. There also exist younger 16-10 Ma basalts and basaltic andesite
flows related to the volcanism of the Snake River Province and the Northern Nevada Rift
40
to the North which include the TR-77 sample above with an age of 14.10 +/- 0.12 Ma
(Gonsior, 2006).
Four ages from the Sulphur Springs Range in central Nevada, east o f the FCM
have been determined and further discussed in Ryskamp et al. (2008). The U-Pb zircon
ages o f these units are as follows; an andesite (31.4 + 1.3/-0.5 Ma), plagioclase dacite
dome (35 +/- 0.5 Ma), biotite dacite tuff (35.5 +/- 0.4 Ma) and biotite porphyry intrusion
(35.9 +/- 0.5 Ma)
A number of tuffaceous units have also been dated from the FCM and Shoshone
ranges presented in McKee et al., (1971). FCM K-Ar and fission track dated samples are
as follows; FCM rhyolitic welded tuff (24.4 Ma), Bates Mountain rhyolite welded tuff
(23.1 +/-1.7), and Caetano rhyolitic welded tuff (31.2 - 33.3 +/- 2.7 Ma). Shoshone
Range K-Ar and fission track dated samples include; Rhyolitic Caetano welded tuff (31.3
Ma), Bates Mountain rhyolitic welded tuff (24.7 +/- 1.0 Ma), and rhyolitic welded Tuff of
McCoy Mine (26.3 +/-1.6 Ma). However, in contrast to the dated tuffaceous units,
Paleogene rhyolitic to basaltic lava flows from the FCM and surrounding areas have been
studied far less. Dates determined by K-Ar method for other Paleogene flows include the
Clan Alpine Mountains Andesitic flow (35.0 +/- 1.2 Ma), and the Simpson Park
Mountain flows; Dacite flow (34.5 Ma), glassy rhyolite flow (30.9 +/- 0.7 Ma) and an
andesite-dacite flow (35.4 Ma) (McKee et al., 1971).
Four tuffs dated from Bates Mountain, which is located southwest o f the FCM,
include the Nine Hill tuff (25.27 Ma), Tuff of Campbell creek (28.6 Ma), Tuff of
Sutcliffe (30.48 Ma), and Tuff of Rattlesnake Canyon (31.03 Ma) (John et al., 2008).
41
4.3 New Ar-Ar ages
O f the 29 Paleogene mafic to felsic samples analyzed in this study, three samples
have been dated using Ar-Ar methods by Chris Henry at the Nevada Bureau o f Mines.
Samples H-10-53 (Rhyolite dome), 10-BV-07 (Andesite) and 10-BV-47 (Dacite), were
each sampled from different geographical areas of the FCM and surrounding area and
yielded ages of 34.24 +/- 0.05 Ma, 33.3 +/- 0.3 Ma and 33.82 +/- 0.14 Ma respectively
(C. Henry, pers. Comm., 2011). Since they all overlap with or close to their uncertainty
ranges we can deduce that the eruption o f these Paleogene lavas seems to have occurred
within a relatively short time period of 1 Ma between 33.3 and 34.3 Ma.
In addition to these dated samples, younger Fish Creek Mountain rhyolite ash
flow tuff (FCMT) has also been dated using sanidine and further discussed in Varve
(2013). The ages and locations are as follows; 10-BV-38 sampled from the western
Golconda Canyon o f the Tobin Range (24.95 +/- 0.08 Ma), 10-BV-17 sampled from the
Horseshoe Basin o f the Fish Creek Mountains (24.91 +/- 0.05 Ma), and H03-73 sampled
from the southern Fish Creek Mountains (24.88 +/- 0.05 Ma).
42
Chapter 5: Petrography of Paleogene volcanic rocks of the FCM
5.1: Introduction
The purpose o f this chapter is to describe the petrographic characteristics o f the
Paleogene lava flows in the FCM. Textural differences were observed in both matrix and
phenocryts/xenocrysts. The mineralogy of samples described in this section is done in
order o f abundance in thin section. The petrographic descriptions of each slide are
presented in Appendix B, and photographs of representative thin section slides are
included in figures 15 to 23. Plagioclase phenocrysts and xenocrysts are discussed in
terms of degrees of disequilibrium; low degree describes phenocrysts with only a slight
appearance o f sieve texture, which is a sign of disequilibrium, moderate degree describes
phenocrysts with abundant sieve texture which looks as if it is eating away at the crystal
from outwards to in (Figure 19 and 20); high degree describes xenocrysts which are
highly sieved and are anhedral with a precipitated rim which is potentially in equilibrium
with the surrounding melt (Figure 15).
5.2 Basaltic Andesites
Basaltic andesites were mostly samples from the Tobin Range within Golconda
Canyon; however one sample (16) is from the eastern Horseshoe Basin. They typically
contain phenocrysts o f plagioclase, pyroxene, olivine, amphibole, biotite and K-feldspar.
In these rocks the abundance o f plagioclase ranges from 15-20% where plagioclase is
present as both the most common phenocrystic and groundmass phase and where
trachytic texture of plagioclase laths is common. Plagioclase phenocrysts vary from 0.5 to
5 mm long and typically show minor seritization. Compared to other rock types,
43
plagioclase phenocrysts seem to show little to no disequilibrium with the melt.
Plagioclase xenocrysts were found in sample 42, indicative of some crustal
contamination, and display high disequilibrium with the melt (Figure 15). Olivine
phenocrysts often show iddingsite weathering along fractures and vary in size from 0.5 to
1 mm (Figure 16). Porphyritic hornblende and biotite (0.5-1.5mm) are more apparent in
sample 40 with reaction rims which represent that they were in disequilibrium with the
melt during crystallization either due to decompression, degassing or assimilation.
Pyroxenes also occur as a phenocrystic phase (0.5-5mm) with no apparent reaction rims.
Opaque minerals vary from moderate to abundant throughout the samples. For the most
part, surrounding the phenocrysts is a glassy to microcrystalline groundmass of very fine
grained feldspars and opaques. Samples 39 and 40 (taken in proximity o f each other) are
vesicular and less fresh where as samples 16 and 42 are the most mafic samples and
fresh. Sample 39 and 40 also contain K-feldspar (2-3%) where as samples 16 and 42 do
not.
5.3 Andesites
Andesites can be divided in terms of trachytic and non-trachytic. Andesites
contain phenocrysts of plagioclase, pyroxene, K-feldspar (Carlsbad twinning), biotite and
hornblende. Alteration is much more common in the andesites in contrast to the basaltic
andesites. Chloritization present in the andesites is often associated with biotite, pyroxene
or amphiboles and sericite with plagioclase and K-feldspar (Figure 17). Plagioclase is
typically euhedral with moderate disequilibrium with the melt, and is usually zoned.
Plagioclase phenocrysts also display sieve texture, indicative o f magma mixing and
disequilibrium, which is very common in the western andesites.
44
Figure 15: Basaltic andesite sample 10-BV-42 shows a plagioclase xenocryst displaying a
high degree of disequilibrium with the melt. It is sieved and almost completely rounded
(anhedral) and has a precipitated rim which is possibly in equilibrium with the
surrounding melt. Its presence is also indicative o f some crustal contamination.
Olv
£
m m
Figure 16: Basaltic andesite sample 10-BV-42 shows olivine phenocrysts with iddingsite
weathering in its fractures varying in size from 1-2 mm. The surrounding groundmass
consists o f glassy very fined grained to microcrystalline composed o f finer grained
phases plus opaques.
45
Figure 17: Highly altered andesite sample 10-BV-05 showing chloritization (light green
in ppl) on the left and seritization o f plagioclase phenocrysts in xpl on the right.
Figure 18: Andesitic sample 10-BV-41 under xpl shows a hornblende opaque reaction
rim indicating some disequilibrium with the melt either due to decompression or
degassing. Plagioclase seems to be in moderate disequilibrium with the melt.
46
Biotite and hornblende show evidence o f being in disequilibrium as they have opaque
reaction rims that sometimes extend into the core. This suggests either shallow degassing,
or slow magma ascent across the amphibole stability field limit (Rutherford et al., 1998)
(Figure 18). Sample 12B is the most altered, which was collected from a volcanic breccia
and has both seritization o f plagioclase and Mg rich chloritization of pyroxenes, biotite
and amphiboles based on its first order interference colours. Samples 10-BV-46A and 10BV-41 contain abundant lithic clasts of limestone, based on the presence o f carbonate
minerals in thin section. Vesicles vary from 1-2%. The groundmass is glassy to
microcrystalline.
The trachyandesites can be examined in terms of vesicular versus non-vesicular.
Commonly both contain a trachytic texture of plagioclase laths in a glassy to
microcrystalline groundmass. The non vesicular samples (7-10,45) were taken from
columnar jointed outcrops typical o f thicker flows and contain phenocrysts of
plagioclase, pyroxenes, olivine, K-feldspar, hornblende and biotite. Megacrysts of
plagioclase are also present in samples 07 and 09 up to 1 cm and also display sieve
texture (Figure 19). Plagioclase minerals are also sometimes zoned with moderate
disequilibrium. Zonation is typically found on the larger plagioclase phenocrysts and
megacrysts, which indicates a possible recharge by magmas from below as the recharged
magmas may have different Ca/Na ratio. Large reaction rims are present on biotite and
hornblende in sample 07 and trace epidote is found in sample 45. Vesicular samples (293 2 ,34A, 49) typically contain phenocrysts o f plagioclase, clinopyroxene, K-feldspar,
olivine, biotite and hornblende. Plagioclase and K-feldspar are often seritized and
sometimes zoned with evidence for moderate disequilibrium. Sample 31 is the most
47
vesicular containing as much as 5% vesicles (Figure 20). There is a lack o f reaction rims
around biotite and hornblende in the vesicular samples.
5.4 Dacites
Dacites, classified as trachy-dacites according to the TAS discrimination diagram,
vary from alkaline (11,34B) to subalkaline (6, 44, 46B, 48A). Alkaline and subalkaline
samples contain porphyritic plagioclase (1-3 mm) with minor biotite and pyroxene
phenocrysts. In both alkaline and subalkaline samples, plagioclase phenocrysts are
sometimes seritized, display sieve texture, are partly resorbed and are thus in moderate
disequilibrium with the melt (Figure 21). Reaction rims around porphyritic plagioclase
are only present in sample 34B whereas biotite reaction rims are more common in the rest
of the samples. Biotite is chloritized in sample 11. Vesicle percent varies, but generally
the dacites contain 1-2 % vesicles for both alkaline and subalkaline types. Subalkaline
sample 46B contains plagioclase megacrysts (12- 15mm) and spherulites are apparent in
thin section under plane polarized light (ppl) (Figure 22). Sample 10-BV-48A contains
rounded xenocrysts of sericitized K-feldspar indicative of crustal contamination. Other
phenocrystic phases in subalkaline dacites include K-feldspar, biotite and pyroxene. The
groundmass is glassy to microcrystalline.
5.5 Rhyolites
All rhyolitic samples are subalkaline. Four out of 29 samples were classified as
rhyolite, 2 o f which rest on the dividing line in between dacites and rhyolites. They
typically contain phenocrysts of plagioclase, K-feldspar, biotite, pyroxene and minor
quartz; however, rhyolites tend to be the least porphyritic out o f all rock types.
48
Figure 19: Andesitic sample 10-BV-09 under xpl displaying a large zoned plagioclase
megacrysts with sieve texture.
Figure 20: Andesitic sample 10-BV-31. This sample is highly vesicular and plagioclase
phenocrysts display sieve texture and seem to be in moderate disequilibrium with the
melt.
49
2 mm
Figure 21: Dacitic sample 10-BV-34B under xpl. Plagioclase phenocrysts are seritized,
show sieve texture, are partly resorbed and display reaction rims in moderate
disequilibrium.
Figure 22: Dacitic sample 10-BV-46B under ppl. Sample shows spherulites (Sph) as
rounded structures which indicates divitrification and plagioclase phenocrysts are in
moderate disequilibrium.
50
Plagioclase phenocrysts typically vary from 1 to 5 mm and are commonly in moderate
disequilibrium with the melt. Zoned plagioclase is common in sample 12A classified as a
volcanic breccia in outcrop. Rounded plagioclase and K-feldspar are also common in
sample 12A, interpreted to be xenocrysts from crustal contamination. Reaction rims are
found mostly around biotite and less so around plagioclase. Spherulites are also present in
plane polarized light of sample 48B as rounded structures (Figure 23). Biotite, like
plagioclase, appears to be in moderate disequilibrium in sample 47 based on opaque
reaction rims. All rhyolites are vesicular and contain a glassy to microcrystalline
groundmass.
5.6 Summary
All rocks types include lavas with <35% phenocrysts, <10% vesicles, although
most samples have <25% phenocrysts. Amongst all rock types, plagioclase exists as both
the most common phenocrystic phase and is often found in the trachytic groundmass as
laths. Microporphyritic is a common texture of the basaltic andesites where most o f the
weakly porphyritic samples exhibited trachytic texture. Typical microscopic features seen
in most samples include sieve texture in phenocrystic plagioclase, oscillatory zoned
plagioclase, amphibole phenocrysts with reaction rims, plagioclase microlites and the
presence of glass in the groundmass of several of the more weakly porphyritic rocks.
These textures either suggest shallow magmatic process or that these flows cooled
extremely rapidly. Other phenocrysts phases typically include pyroxene (5-10%), Kfeldspar (1-10%), biotite (1-8%), olivine (1-5%), hornblende (1-3%), and quartz (1-2%)
with most phenocrysts being euhedral to subhedral. Phenocrysts range from millimeters
to up to 1.5 cm; plagioclase and hornblende are commonly the largest. In andesitic and
51
dacitic rock types, reaction rims are commonly found around biotite and hornblende in
varying intensities either the result of decompression and/or degassing and H 2O loss from
the magma. In terms of alteration, biotite is commonly altered to chlorite, plagioclase and
K-feldspar to sericite, and olivine to iddingsite. The visible matrix grains are notably
small compared to the phenocrysts and consists o f the same phases as are present as
phenocrysts, as well as glass and opaque’s. Vesicles are common in all rock suites,
especially in eastern Horseshoe Basin samples, except for samples taken from columnar
jointed outcrops typical of thicker flows.
52
2 m m
Figure 23: Rhyolitic sample 10-BV-48B under ppl. The rounded structures and
spherulites (sph) and the elongated brown-black minerals are biotite with reaction rims in
a glassy microcrystalline groundmass.
53
Chapter 6: Geochemistry; Geochemical and Radiogenic Isotope Svstematics of
Paleogene felsic to mafic rocks of the Fish Creek Mountains, north-central Nevada.
Western United States.
6.1: Introduction
Whole rock, trace element and isotopic data from the study area are presented in
Appendices C, D and E. The purpose o f this chapter is to use the geochemical data to
describe the chemical characteristics of these rocks. For plotting all major and trace
element analyses have been recalculated on a volatile-free basis. Geochemical data from
29 collected rock samples are plotted on the conventional total alkalis silica (TAS) plot.
O f the 29 samples analyzed, 4 are classified as basaltic andesites, 15 are andesites, 6 are
dacites, and 4 are rhyolites (Figure 24; Le Bas et al., 1986). With the exception o f 4
samples that are slightly alkaline, the samples are sub-alkaline. On the AFM diagram
(Figure 25) the samples mostly plot in the calc-alkaline field o f Irvine and Baragar
(1971). Magnesium (Mg) numbers (Mg#=100x (Mg/Mg+Fe2+)) where Fe2+ = 0.9 Fe
total) range from about 13-56 with an average of 40. Whereas the basaltic andesites have
restricted Mg#’s in between 48 and 57, the andesites display much more scatter with
values between 30 and 55.
6.2 Whole rock major geochemistry
The volatile free major element data from all 29 samples are plotted on Si02
variation diagrams (Figure 26). The purpose of these diagrams is to present the data
visually and to display the variation in abundance of the elements in the 4 rock types
found in each o f the areas within the Fish Creek Mountains described above.
54
in
I Tertiary mafic-felsic
Phortolite
Foidite
Trachyte
Trachydacite
o
Phonotephrlte
Rhyolite
Basaltic'
Tephrite
Basanita
basalt
in
o
4 0
5 0
6 0
7 0
8 0
Si02 (Wt%)
Figure 24: Total alkalies vs. silica diagram o f Le Bas et al. (1986), using recalculated
analyses all in wt (%). Dotted line represents the division in between alkaline and
subalkaline from Irvine and Baragar (1971).
55
■ Basaltic andesite
• Andesite
A Dacite
♦ Rhyolite
Tholeiite S eries
Calc-alkaline S e rie s
MgO
Figure 25: Paleogene mafic to felsic samples plotted on a MgO, FeO1, and Na20+K20
(wt%) ternary diagram (Irvine and Baragar 1971).
56
Some of the major element oxides, particularly the alkalies (K 2O and Na20 ) are known to
be mobile during alteration (e.g., Hughes 1973). As a result, major element oxide
discrimination diagrams that use these elements are likely to be less reliable than
diagrams involving the less mobile, high-field strength elements such as Ti, Zr, Nb and Y
(Winchester and Floyd 1977). Hence the latter elements carry more weight in the
interpretations presented here.
Variations in the major element chemistry, taking into account the observed
phenocrysts phases can help constrain the differentiation history of these volcanic rocks.
The decrease of FeO* and T i02 with increasing S i02 and the presence o f opaque minerals
in most rock samples, especially the most mafic samples, illustrate the importance of
titano-magnetite as a fractionating phase. MnO vs. Si02 diagram shows a similar
tendency to the Fe plot which reflects the fact that Mn readily substitutes for Fe in ironbearing minerals (Figure 26).
A12C>3 varies widely amongst rock types with the exception of rhyolites and
consistent with the abundance o f plagioclase in thin section (Figure 26). The decrease o f
CaO with increased Si02, and small negative Eu anomaly present in more evolved
samples (especially the most evolved dacites and rhyolites) indicate the fractionation of
plagioclase and possibly clinopyroxene, which is the major phenocrystic phase found in
all rock types (Figure 26).
The wide range of Na20 concentrations may reflect sodium mobility and the
differences in the extent o f albitization among the samples. The subtle decrease of P2Os
with respect to increasing Si02 reflects the fractionation o f apatite (Figure 26).
57
*•
■ * *
£
■
A * A
0
1
A
A
a
•
*
•A
SO
55
60
65
50
70
SS
60
so*
!
V
66
50
70
55
60
65
70
65
70
S*02
s«j
'A
..............................
A Jkf
'• V
~
* '
3O
A*
▲ A
A
50
58
80
'
65
•••%
♦
70
SO
56
60
SO.
..................................................................
5. *
u
*
* s »
a a
80
SiO*
60
£
fo
a
0T”
0
O
65
<£ *0
*
v
7D
3
3
* 1 t 4
s0
... ... A
55
55
SiOz
8
0
5
50
SO
70
s*0j
wm
■
65
♦4
4
0
50
55
AA
**
♦
66
85
♦;
70
,
3
j
3 1
50
SlOa
Figure 2 6 : Silica variation diagrams for T i C > 2 , A I 2 O 3 , MgO, CaO, N
MnO and Ca0 /Al2 0 3 all in Wt (%). Symbols are as in Figure 25.
58
^
A
.......... *
85
\
a * .♦:
60
S02
a 2 0 , P 2 O 5,
FeOl,
<
jd
High-K calc-alkaline
Series
O
TholeiHe Series
4 5
5 0
5 5
6 0
6 5
7 0
7 5
S 1Q2
Figure 27: K2O vs. SiC>2 plot of Peccerillo and Taylor (1976) where the majority o f FCM
samples plot in the high-K field to shoshonite series with the exception o f one basaltic
andesite that plots in the med-K field (10-BV-42). Symbols are as in Figure 25.
59
When samples were plotted on a K.20% vs. SiC>2% graph from Peccerillo and
Taylor, 1976 (Figure 27), the majority of the samples plotted in the high-K field to
shoshonite series. However, there is one sample (sample 10-BV-42), a basaltic andesite
that plots in the med-K field. The more mafic samples generally have lower
concentrations of K2O relative to samples with high silica concentrations. The highest
values were mostly observed in southern flow rocks, followed by eastern and western
areas.
6.3 Whole rock trace element geochemistry
Combined with major element trends, trace elements can also be used to help
constrain the differentiation history. Similar to A
I 2O 3,
Sr also varies widely amongst rock
types consistent with the variable amount o f plagioclase in most samples where Sr often
substitutes for Ca in plagioclase (Figure 28).
The observed decreases in Ca0 /Al2 0 3 , CaO, and Sc with increasing SiC>2 are
consistent with clinopyroxene fractionation. The decrease of MgO and Co with respect to
Si02 in the more mafic rocks probably reflects olivine fractionation (Figure 28).
Combined with the decreases o f FeO1and Ti02 with increasing Si02, V also
supports titano-magnetite as a fractionating phase. When plotting V vs. Fe 2C>3 in Figure
29, a positive correlation exists which reflects the ability o f V to substitute for Ti in
titanium iron oxides. Cr and Ni are also compatible with ferromagnesium minerals and
are therefore relatively abundant in most mafic rocks, especially samples 10-BV-42 and
10-BV-16 where as the more evolved southern samples tend to have the lowest Ni, Cr
and V.
60
The occurrence of xenocrysts and the large range in Pb, Sr and Nd concentrations
in most rock types regardless of silica content, suggest that crustal contamination may
also be a process affecting the chemistry (Figure 28). Th concentrations may be derived
from the continental crust and display the widest ranges within the dacites with sample
10-BV-l 1 containing as much as 50 ppm Th (Figure 28). The positive correlation that Pb
and Th have with SiC>2 increasing by a factor of five indicates that FCM samples
experienced some Pb and Th contamination in contrast to other incompatible elements
such as Ba or Zr that vary irregularly.
Turning to incompatible element diagrams, there exists a wide range of fluid
mobile (large ion lithophile elements or LELE) to fluid immobile (high field strength
elements or HFSE and REE) elements in these rock types. The incompatible element
diagrams in Figure 30a-d are normalized to primitive mantle after Sun and McDonough
(1988) and are strongly enriched in incompatible elements (e.g. Rb, Ba, Th, U) relative to
primitive mantle and are relatively homogeneous. Negative troughs occur at Nb and Ti in
all samples. Peaks occur in all samples at Pb and most at Sr and K (LILE). Therefore,
fluid mobile elements such as Rb, Ba, U, K and Sr are enriched compared to less fluid
mobile elements such as Nb, Ta, Zr, Ti and the REE (Figure 30a). The incompatible
element diagrams for the basaltic andesites and andesites are relatively homogeneous
with the exception of two samples from the basaltic andesites (10-BV-16 and 10-BV-42)
and two from the andesites (10-BV-05 and 10-BV-41). Basaltic andesite sample 10-BV16 shows an overall higher abundance o f most incompatible elements, especially Pb,
whereas 10-BV-42 shows an overall depletion of incompatible elements whereas still
reflecting the same pattern as seen in the other basaltic andesites (Figure 30b).
61
25
*
65
SO
70
IS
20
55
60
50
70
■ ...............................
■ .................................
TO
65
70
Pb
30
V
100
SO
65
50
70
55
60
65
50
70
55
60
Si02
’
!
*
♦
20
0M
30
40
2000
Ba
1500
300
......................... M ......
200
SiOa
<2
Th
60
S 02
\
10
100
♦
■
1000
• ■ ... * .............
SO
5S
eo
65
70
50
55
60
65
\
Zr
65
10
20
................................♦
55
60
50
■
50
55
SiOj
SiOj
200
s
♦
65
70
60
s«o2
30
40
55
10
Co
1
■
50
♦
s
>
■
10
.
5 *
A S
s
Sc
%
8
•
°
■ ..............................
♦
... ■ .......... ▲ ........
50
70
55
60
■■
r— --------x
.............
1
z
*
•
A
■»
I *
50
55
60
65
70
50
55
60
SiOz
SIO2
65
70
50
55
j
60
65
StOj
Figure 28: Silica variation (wt%) diagrams for Ce, Sr, Sc, Co, V, Pb, Zr, Ba, Th, Cr, Ni
and Nb (all in ppm). Symbols are as in Figure 25.
62
70
&O2
SI02
SIO2
65
70
o
o
CM
o
o
••
o
in
♦ ♦
2
6
4
8
FeaOst
Figure 29: Plot of total iron as Fe2(>3 vs. V. Symbols are as in Figure 25.
63
+-<
c
(0
5
HE?
£
c
CL
■ Basaltic andesite
• Andesite
▲ Dacite
♦ Rhyolite
8
O
a.
E
co
CO
L*
P
r
P
t
o Sr
Lm
n>
c.
C*
R
b T
h
em
Nd
8m
N
O
am
Oy
Yto
Y L
u
0>
73
C
CO
2
0
>
73
8
I
*C
a.
CL
E
CO
w
T
h
Nt>
S
r
•n
Y
Lu
Figure 30: Incompatible element diagrams for (a) all samples and (b) basaltic andesites.
All values are normalized to primitive mantle (Sim and McDonough, 1988). Ta (not
shown) behaves the exact same range as Nb. The lowest sample of the basaltic andesite
plot is sample 42.
64
©
c
©
2
©
>
*4«#
E
*c
CL
jS
CL
E
©
w
©
33
c
©
§
©
.2:
-4—»
I
*C
CL
j !
a.
8
E
©
£/>
LM
p»
Sr
Nd
Ti
Y
Lu
Figure 30: Incompatible element diagrams (c) andesites and (d) dacites and rhyolites. All
values are normalized to primitive mantle (Sun and McDonough, 1988). The lowest
sample of the andesite plot is sample 10-BV-41.
65
The andesite sample 10-BV-05 shows depletion in barium compared to the
overall enrichment and a more prominent depletion of phosphorous, and sample 10-BV41 shows an overall slight depletion of all incompatible elements compared to the rest of
the andesites (Figure 30c). Dacites show an overall enrichment in LREE compared to the
rest of the rock compositions and sample 34B is strongly enriched in Pb, possibly as a
result of crustal contamination. Lastly, the rhyolites plot at the lowest concentrations of
Sr and P, possibly as a result of minor fractional crystallization due to apaptite and some
plagioclase fractionation (Figure 30d).
Trace element ratio plots such as Zr/Nb vs. Ce/Pb (Figure 31) show a pronounced
overlap among basaltic andesites and andesites. Although trends are evident in some ratio
plots when reviewing the data as a whole, grouping among individual rock types is less
common. The plot of Ce/Pb vs. Cs/Rb (Figure 32) distinguish subduction-related fluid
addition characteristics (high Pb) from mantle characteristics (high Ce). The common
features of these plots are that the basaltic andesites are located at both extremes of the
plots as well as in between them.
6.4 Isotope Geochemistry
The initial 87Sr/86Sr of the basaltic andesites range from 0.705907 to 0.707873.
The eNd34 values calculated at 34 Ma of the basaltic andesites range from -1.9 to -6.5.
The andesites range in 87Sr/86Sr from 0.705833 to 0.707098, where as their eNd 34 values
range from -2.5 to -5 (Figure 33a). Overall, an increase in 87Sr/86Sr does correlate with an
increase in silica content (Figure 33c), whereas a very slight negative correlation can be
found when looking at the 143Nd/,44Nd vs. Si02 plot (Figure 33e). Overall, the Sr and Nd
66
isotopic values overlap between all rock groups, however the dacites and rhyolites do
show a more restrictive range compared to the basaltic andesites and andesites. When
87
plotting eNd34 versus Sr/ Sr initial on Figure 33a, the isotopes shows a negative
correlation and the basaltic andesites plot at both high Sr and low Nd, and at high Nd and
low Sr where as more evolved rock types plot in between.
Pb isotopic ratios also plot similarly to the Sr and Nd isotopic ratios. Individually,
the basaltic andesites plot at both high and low 206Pb/204Pb values where as other rock
types plot in between them. Basaltic andesites range in 206Pb/204Pb from 19.25 to 19.40
where as the andesites range from 19.32 to 19.4 (Figure 33b). Also shown in Figure 33d
on the 206Pb/204Pb vs. 208Pb/204Pb plot is a slight positive correlation with an observable
overlap amongst the more evolved rock types.
67
MORB
<g
a
.
M an tle
Inc
Slab fluid
com ponent
M odem S o u th C a s c a d e
B a sa lts
Inc % o f m eltin g
—r~
~ r~
I
I
10
15
20
25
30
Zr/Nb
Figure 31: Trace element ratio plot of Ce/Pb vs. Zr/Nb. X-axis indicates an increase of
partial melting percent and Y-axis indicates an increase of mantle or decrease of slab
fluid components. Modem South Cascades Basalts data from Borg (1997) and BV
(Buffalo Valley) data from B. Cousens pers. Comm. 2013. Symbols are as in Figure 25.
in c M antle
C a s c a d ia s e d s
•
•
CL
e
O
A
*
A
inc Fluid
inc of S e d im e n ts
T
0.00
0.02
0 .0 4
0.06
0.08
C s/R b
Figure 32: Trace element plot o f Ce/Pb vs. Cs/Rb. X-axis indicates an increase of
sediments and Y-axis indicates an increase o f mantle or increase of fluid components.
Cascadia sediments from Prytulak et al. 2006. Symbols are as in Figure 25.
68
*
*8
i t
i ’
s
• ♦
1920
19.25
1920
1920
19.25
19.30
19.35
1940
1945
o
e
d
I©
? Io
?•
©
Q
66
00
95
70
1935
1940
* * P ti/* >Vb
o
■ Basaltic Andesite
• Andesite
A Dadte
♦ Rhyolite
*©
»
A
•a
S
3
«>
d
R
d
80
66
Figure 33: Initial Nd, Sr and Pb isotopic ratios in the FCM and surrounding areas (a, b,
d), (c) 87Sr/86Sr vs. SiC>2 and (e) 143Nd/,44Nd vs. SiC>2 content.
69
C hapter 7: Discussion
7.1 Introduction
This chapter includes discussion of the source of the melt, defines the parental
magma, determines the degree of crustal contamination and defines the rock samples
based on their petrography and geochemistry in an attempt to determine the geotectonic
setting for the emplacement of the Paleogene mafic to felsic samples o f the Fish Creek
Mountains and the neighbouring Tobin and Shoshone Ranges.
7.2 Petrological comparisons
Calc-alkaline basaltic andesites and andesites are the most abundant rock types
from this area and are common throughout the Great Basin, which includes widespread
calc-alkaline rocks o f intermediate to silicic composition capped locally by younger
basaltic and rhyolitic rocks (Best and Christiansen, 1991). The calc-alkaline samples
from the FCM show similarities in both the petrographic and geochemical signatures to
the calc-alkaline samples found in the Sulphur Springs Range in central Nevada
(Ryskamp et al., 2008), the North Clan Alpine and Stillwater Range o f central Nevada
(A. Timmermans, pers. comm. 2012), Western Great Basin (WGB) (Ormerod et al.,
1988,1991) and study areas of the Ancestral Cascades Arc (Cousens et al., 2008; Stoffers
2010; Clark 2011).
Petrographic features in common with all areas include sieve texture in
phenocrystic plagioclase, oscillatory zoned and partly re-sorbed plagioclase, non­
equilibrated hornblende and biotite phenocrysts, plagioclase microlites and the presence
o f glass in the groundmass. Besides abundant plagioclase phenocrysts, generally other
70
phases include pyroxenes, K-feldspar, biotite, hornblende, olivine (intermediate samples),
and quartz (felsic samples) in decreasing abundances with most phenocrysts are euhedral
to subhedral ranging from millimeters to up to 1.5 cm.
Petrographically, the FCM lava flows compare favourably to the MiocenePliocene porphyritic lava suite o f the Ancestral Cascades defined by Cousens et al.
(2008). The porphyritic suite is characterized by abundant, highly zoned, blocky
plagioclase megacrysts as large as
1
cm which are commonly reverse or oscillatory zoned
and also commonly have sieve texture. This zoning and texture are also commonly found
in FCM andesites (Figure 19). Euhedral to subhedral pyroxenes are the next most
abundant phenocrysts in both the FCM and Ancestral Cascades typically making up 515% of the rock. Amphibole crystals range in size from 1-5 mm and constitute 1-10% of
the rock (Cousens et al., 2008). Typically, amphibole crystals have opaque reaction rims
that may extend to the core of the grain, where the hornblende has broken down to an
anhydrous assemblage of plagioclase, pyroxene, and magnetite. This texture is also found
in FCM andesite and dacite samples (Figure 18). Biotite occurs as subhedral phenocrysts
(1-3 mm) and can be deep red in thin section due to oxidation and iron oxide inclusions
(Cousens et al., 2008). Biotite grains present in FCM samples are typically this colour;
however, also contain opaque reaction rims commonly found around amphiboles. The
porphyritic lava suite groundmass is dominated by plagioclase, accompanied by Fe-Ti
oxides and high SiC>2 glass, typically the common groundmass found in all rock types o f
the FCM.
71
The most mafic basaltic andesites of the FCM typically consist o f pyroxene,
olivine and some plagioclase with a matrix o f plagioclase, Fe-Ti oxides, pyroxene and
minor glass. Trachytic flow texture is common and defined by elongate plagioclase
phenocrysts all of which is commonly also found in the basaltic andesite dikes o f the
Sulphur Springs Range (Ryskamp et al., 2008). Calc alkaline basalts of the N. Sierra
(Staffers, 2010) are somewhat comparable to the FCM basaltic andesites as well,
however N. Sierra basaltic andesites contain greater abundances of olivine and fewer
pyroxene phenocrysts, much like the alkali olivine basalts o f the WGB o f Ormerod et al.
(1988).
Andesitic flows of the Sulphur Springs and North Clan Alpine and Stillwater
Ranges demonstrate textures disequilibrium textures that could be indicative of magma
mixing, including sieve texture o f phenocrystic plagioclase, re-sorbed felsic phases, thick
reaction rims and coexisting mafic olivine with felsic quartz (Ryskamp et al., 2008; A.
Timmermans, pers. comm. 2012). The Sulphur Springs andesite flows; however, contain
unstrained quartz and sanidine megacrysts along with olivine and pyroxene which are not
present together in the andesitic rocks of the FCM. The quartz phenocrysts are also
extensively resorbed and many have reaction rims of clinopyroxene (Ryskamp et al.,
2008). Sieve texture in phenocrystic plagioclase, which is present in andesites (Ryskamp
et al (2008), A. Timmermans (pers. comm. 2012), Cousens et al (2008) and Ormerod et al
(1988, 1991)), is indicative of non-equilibrium crystallization or re-melting and was
likely triggered by magma mixing within the system which is very typical o f Western
Great Basin andesites (Cousens et al., 2008). Resorption is observed predominately in
plagioclase. Amphibole and biotite show evidence of being in disequilibrium as they
72
have thick opaque reaction rims, suggesting that either the melt degassed beneath the
volcano or the magma ascended slowly across the amphibole and biotite stability fields
(Rutherford et al., 1998). Andesitic sample 10-BV-08 contains about 5% olivine, atypical
for andesitic lavas of the western USA and could represent mixing of a more mafic
magma (containing olivine and pyroxene) with a more felsic magma (containing Kfeldspar). Zonation was also commonly found on many of the large plagioclase
phenocrysts indicating possible recharge by magmas from below as the recharged
magmas may have different Ca/Na ratios.
The dacitic lava domes, basaltic andesite dikes, andesite dikes and lava flows, and
flow banded rhyolite flows of the Sulphur Springs Range of central Nevada are the most
petrographically comparable to the samples of the FCM. The Dacitic lava domes
resemble the dacitic-rhyolite dome of sample 10-BV-48 (Figure 12) sampled from the
Shoshone Range. They both have distinctively large plagioclase phenocrysts with lesser
amounts of quartz, K-feldspar and pyroxene, along with amphibole, biotite, and Fe-Ti
oxides. Alteration is present in both types including chlorite, clay minerals, and iron
stains where biotite is commonly altered to chlorite and feldspars to sericite (Ryskamp et
al., 2008). Flow banded rhyolites of East Sulphur Springs are similar to the crystal poor
rhyolites of the FCM. They both contain extremely small phenocrysts o f quartz and
plagioclase in a glassy matrix (Ryskamp et al., 2008).
Chlorite alteration is found throughout most samples as an alteration product of
mafic minerals such as pyroxene and amphibole; however it is most commonly found
around biotite. In this environment chlorite may be present as a metasomatism product
via the addition of Fe, Mg or other compounds into the rock mass. In thin section,
73
chlorite is predominately Mg-rich based on its first order interference colours in contrast
to anomalous blue in the case of Fe-rich chlorite. Sericitic alteration often found on
plagioclase phenocrysts implies acidic conditions which may be associated by the
passage o f hydrothermal fluids.
Whereas the petrography may be insightful regarding crustal residency time of the
magmas it cannot be relied upon for a means of determining the source of the basaltic
andesites found in the FCM.
7.3 Geochemical comparisons
By analyzing geochemical trends we can begin to deduce what the main lava type
is within the FCM. As noted in chapter 6 , lavas from the FCM and surrounding Tobin
and Shoshone Ranges are all calc-alkaline and generally more evolved than Paleogene
lavas from the Sierra Nevada (Cousens et al., 2008; Staffers, 2010, Clark, 2011), and
most resemble the andesite, dacite and rhyolite flows from the Sulphur Springs Range in
central Nevada (Ryskamp et al., 2008).
The majority of the FCM samples are all highly potassic plotting in the high-K to
shoshonite series on the K20 vs. S i0 2 diagram of Peccerillo and Taylor (1976) (Figure
34). FCM samples are more potassic than the Sierra Nevada samples (Staffers 2010 and
Cousens et al., 2008) and generally plot in similar ranges as the lava flows from the
Sulphur Springs Range in central Nevada. Calc-alkaline affinity, known to be widespread
in the central Great Basin (Best et al., 1989) is shown on tectonic discrimination Figure
35 from Wood (1980) where all of FCM, Sulphur Springs and Ancestral Cascades
samples plot within the CAB field.
74
■ FCM
» Sulphur Springs
▲ Ancestral C ascad es
Shoshonite Series
High-K calc-alkaline
Series
«
<N
Tholeiite Series
o
50
4 5
6 0
5 5
6 5
70
7 5
SiCfe
Figure 34: K2O vs. SiC>2 diagram of Peccerillo and Taylor (1976). FCM samples plot
higher than the Ancestral Cascades into similar and higher potassic values than the
Sulphur Springs.
H03
■ FCM
♦ Sulphur Springs
A Ancestral Cascades
mors
IAT
ni
Ta
Figure 35: Ternary diagram of Th, H f and Ta discriminating between calc-alkalic basalts
(CAB), within plate alkali (WPA), within plate tholeiitic (WPT) and E-MORB, N-MORB
and island arc tholeiitic (IAT) (Wood, 1980).
75
FCM samples exhibit relatively low Mg #’s (48-57 for basaltic andesites)
consistent with the Sulphur Springs area. Both volcanic suites have compositions that are
generally consistent with a subduction zone origin including low Mg # ’s, high K2 O, and
similar “spiky” trace element patterns with LILE enrichment and HFSE depletion
(Ryskamp et al., 2008). Both regions were also affected by the same Paleogene
detachment or roll back of the Farallon plate, which contributed to continental arc
magmatism over a wide area of western North America (Ryskamp et al., 2008).
The chemical variations of all rock types are consistent with at least some
fractional crystallization of the observed phases; feldspars and mafic silicates and oxides.
The FCM volcanic rocks range in composition from olivine bearing basaltic andesite (5%
MgO, 112 ppm Ni, and 263 ppm Cr) to rhyolite (with <1% MgO, 3 ppm Ni, and 29 ppm
Cr). These steep declines in compatible element concentrations suggest fractional
crystallization of mafic mineral phases that have high partition coefficients for these
elements. Other compatible elements, such as Ti0 2 , Fe2 0 3 , MgO, CaO, Sc, and V,
decrease in concentration as Si0 2 increases (Figure 26). Therefore, fractional
crystallization appears to have played a role in the evolution o f FCM samples; however,
the lack of significant Eu anomaly (Figure 44 and 47) combined with the relative
homogeneous incompatible element patterns from basaltic andesite to rhyolite suggest
fractional crystallization was not the dominant process in forming magmas o f evolved
compositions.
Sr-Nd-Pb isotopic studies consistently show that intermediate to silicic rocks
contain large proportions o f ancient continental crust represented by high
low
144Nd/143Nd
87Sr/86Sr
and
ratios (Figure 33a). When comparing FCM radiogenic isotopic ratios o f
76
Sr and Nd to the WGB suites and Ancestral Cascades, the FCM flows cannot be related
by crystal fractionation due to their large Sr isotopic variance and the difference in
incompatible element patterns among the three locations. Crystal fractionation of a melt
will not alter the isotopic signature of the residual liquid because isotopes cannot
fractionate, yet it can increase the overall incompatible element abundances of the liquid
as the silica content increases (Figure 36).
In terms of incompatible elements, all samples have negative Nb anomalies and
high concentrations of LILE relative to HFSE, all characteristic of calc-alkaline basalt of
the central Great Basin (Figure 30) (Best et al., 1989). This suggests a calc-alkaline
magma as our primary lava end member. The enrichment or peaks of the mobile elements
Rb, U, K, Pb and Sr compared to the depletions or troughs of the immobile elements
(HFSE) could reflect melts of the mantle sources that are metasomatized by hydrous
fluids from the subducting slab, as hydrous fluids released off the slab would be enriched
in the mobile elements relative to the immobile elements (Best et al., 1989).
Middle to heavy REE patterns do not show a steep garnet stability field pattern
and instead show a relatively flat pattern with no significant depletions suggesting that
they were generated at depths within the spinel peridotite field. Garnet is known to retain
the HREE in its chemical structure so we would see an appreciable depletion in HREE if
the melts were generated within the garnet lherzolite field (Borg et al., 1997) (Figure 30).
In addition, Cr abundances with the basaltic andesites suggests that a spinel peridotite is
the likely source of partial melts given the compatible nature o f Cr in garnet.
77
0.5132
0.513
a
0.5128
JSontaM O R B
Modem
South v .
Cascades
3?
? 0.5126
Ancestral
Cascades
This
Study
4
0.5124
0.5122
0.7020
Crust4
0.7030
0.7040
0.7050
0.7060
0.7070
0.7080
87 Sr / 86(Sr
Figure 36: Radiogenic isotopic 87Sr/86Sr vs. 143Nd/144Nd plot showing isotopic Sr and Nd
ranges for known studies in the Great Basin. FCM samples are shown and plot within the
yellow field, BV = Buffalo Valley, and WGB = Western Great Basin. BV data from
Wetmore (2011). Modified from Cousens et al. (2008).
78
Lavas from the Ancestral Cascades area (N. Sierra, Lassen Volcanic Centre and the Lake
Taheo-Reno area) and Sulphur Springs (Ryskamp et al., 2008) also display no noticeable
HREE depletions suggesting that they are also melts from the spinel lherzolite field
(Staffers, 2010) (Figure 39).
7.4 Sources of the Parental Melts
Let us first evaluate the possible source of partial melts that could potentially
generate the basaltic andesites in the FCM. They are 1) intraplate mantle, 2) mantle
wedge, and 3) lithospheric mantle.
Basalts erupted in an intraplate setting, such as late Cenozoic Mojave Desert
basalts, are enriched in most incompatible elements and LILE, do not have a Nb
anomaly, and have REE patterns decreasing from light to heavy when compared to
primitive mantle, consistent with a garnet-bearing lherzolite source (Pearce, 1983;
Schmidt et al., 2008). Intraplate Mojave Desert basalts also have higher Nd and lower Sr
radiogenic isotopic ratios, similar to those o f the Buffalo Valley Pliocene - Quaternary
cones found in the FCM (Wetmore, 2011) (Figure 36). The Buffalo Valley cinder cones,
located on the northwestern margin o f the FCM, are thought to be derived from low
degree partial melting of the upper asthenospheric mantle, that also erupted as intraplate
basalts (Wetmore, 2011). The lack o f an Nb anomaly in both the Mojave Desert and
Buffalo Valley cinder cones, consistency with a garnet-bearing lherzolite source in
contrast to spinel lherzolite combined with isotopically lower Sr and higher Nd ratios
relative to the higher Sr and lower Nd ratios o f the FCM, suggest that the FCM lavas
79
were not derived from an intraplate setting like that of the Mojave Desert basalts and
Buffalo Valley cinder cones.
Fluid metasomatized peridotites of the mantle wedge have been proposed to be
the main sources of modem Cascade Range melts such as those at Lassen Volcanic centre
in the Modem South Cascades (Borg et al., 1997) and in the Ancestral Cascades Arc
(Cousens et al., 2008). Figure 36 shows that mantle wedge sources like the Modem South
Cascades plot at high ,43Nd/i44Nd and low
87Sr/86Sr
ratios. Whereas the Ancestral
Cascades also plot in the mantle wedge area, they also extend and overlap into
lithospheric mantle sources indicating that the source of the Ancestral Cascades contains
both mantle wedge and lithospheric mantle components (Cousens et al., 2008). Borg et
al., (1997, 2002) showed that Lassen region calc-alkaline basalts can be defined by the
degree of S r enrichment over other incompatible elements.
S r /P p n m
(primitive mantle
normalized) can be used instead o f LILE/REEs to show the relative enrichment of Sr over
middle REE in arc rocks due to Sr addition via fluids from the subducting slab at the time
of melting (Borg et al., 1997). Variations in Sr and P are not simply due to fractional
crystallization o f plagioclase and apatite which can be seen by the lack o f Eu or P
anomalies in intermediate FCM samples. In Figure 37, FCM basaltic andesites plot in
similar ranges as the WGB basalts and Cascadia sediments but at relatively higher initial
Sr isotopic ratios and the same relatively low Sr/P values. Borg et al. (1997) state that low
Sr/P and high Sr ratios indicates a source from a lithospheric mantle whereas high Sr/P
and low Sr ratios indicate a mantle wedge source (South Cascades).
80
0.7065
Basalts Only
0.70600.7055-
■
FCM
A .
Ancestral C ascades
0.7050CO
S
0.7045-
tf?
0.7040-/
00
0.7035-
0.702
MORB
S r / P pmn
Figure 37: Primitive mantle normalized plot o f Sr/Ppnm versus 87Sr/86Sr o f FCM samples
compared WGB and Sierra Nevada field areas. Block A represents Precambrian
lithopheric mantle and block B represents mantle wedge with a high fluid component.
The increase of 87Sr/86Sr indicates an increase of lithospheric mantle age. Modified from
Cousens et al., 2008.
81
Borg et al. (2002) also state that a high Sr isotope ratio and low Sr/P component is not a
fluid component and must either be an enriched component in the mantle wedge or a
87
8ft
lithosphere mantle component. The increase of Sr/ Sr also indicates tin increase o f
lithospheric mantle age. In contrast, the Ancestral Cascades plot at mid ranges for both
Sr/P and 87Sr/86Sr along a mixing line between lithospheric mantle (A) and mantle wedge
(B) suggesting a source from both components; however, at lower 87Sr/86Sr ratios
suggesting a younger lithosphere component such as Phanerozoic mantle lithosphere
(Figure 37; Cousens et al., 2008). The fact that FCM samples plot at much more
restricted Sr and Nd isotopic values characteristic of the lithosphere and crust compared
to the Modem and Ancestral Cascades combined with lower Sr/P values and higher
>0.706 initial 87Sr/86Sr suggests that FCM sources are not derived from the mantle wedge
like those of the Modem South Cascades and a component in the Ancestral Cascades.
Instead they are inferred to be derived from enriched components in a ca. 1 Ga
Precambrian lithospheric mantle (Figure 37).
The alkali olivine basalts of the WGB do show some similarities to the lavas of
the FCM. WGB basalt incompatible element patterns include subduction signatures much
like those of the FCM. WGB mantle sources also include low Nd ratios and high initial
Sr ratios compared to the Ancestral and Modem Cascades, plotting in similar isotopic
ranges of the FCM on the
144Nd/,43Nd
vs.
87Sr/86Sr ratio
plot (Figure 36). The WGB has a
distinctive subcontinental lithospheric mantle source which was metasomatized by earlier
subduction episodes beneath the western margin of the south-western U.S. (Ormerod et
al., 1991). Thus magmas derived from this source have incompatible element patterns
that resemble modem subduction-related magmas but have much higher 87Sr/86Sr and
82
lower I43Nd/,44Nd that reflect the ancient age of the lithosphere much like isotopic ratios
found in the FCM. In western North America, the western margin of Precambrian North
America is most clearly resolved along the western Idaho Shear zone which is a
subvertical mylonitic structural boundary coincident with the abrupt change in initial
87Sr/86Sr
of Mesozoic and Cenozoic magmatic rocks (Figure 38). Here, initial 87Sr/86Sr
increases rapidly from <0.706 in the west to >0.706 in the east across the western Idaho
shear zone (Armstrong et al., 1977; Manduca et al., 1992). Continental material west of
the cratonic margin in westernmost Idaho and central Nevada consists o f PaleozoicMesozoic oceanic volcanic arcs, accretionary prism complexes and associated basinal
successions accreted to western North America during Middle to Late Jurassic arccontinent collision (Dickinson, 1979). In terms of subduction and primitive mantle
signatures (Figure 37), the FCM lava flows are very similar to the WGB enriched
lithospheric melts which suggests the primitive mantle source of the FCM has also been
metasomatized by earlier subduction episodes. However, in Figure 36, the FCM samples
are more restricted and more “crustal”, crossing to the east side of the 0.706 initial Sr line
that defines the western margin of Precambrian lithospheric mantle, they are calc-alkaline
high-K lavas compared to lower-K WGB alkali lavas, and show evidence o f larger
volcanic edifices with rare dome collapse and debris flow deposits that are lacking in the
WGB (Ormerod et al., 1991).
In terms of REE patterns, the FCM suite and the Sulphur Springs Range are
similar to igneous rocks from continental margin subduction zones (Ewart, 1979) with
relatively steep slopes and small negative Eu anomalies. Like volcanic rocks erupted in
83
other subduction settings, all of the rocks have negative Nb and Ti anomalies and positive
anomalies for Pb and other LILE on primitive-mantle normalized diagrams (Figure 39).
According to Figure 40 (basaltic andesites only), the FCM and Ancestral Cascades both
follow similar incompatible element patterns, hinting at similar origin; however overall
the FCM flows are more enriched in incompatible elements and the Ancestral Cascades
have greater Nb, Ti anomalies and higher MREE to LREE (Sr/P) ratios that represent a
mantle wedge component that is not present in FCM samples. Sample 10-BV-42 seems to
be more depleted and plots at similar concentrations to the Ancestral Cascades. Variance
in the incompatible element patterns for the calc-alkaline samples may be the result of: 1 )
the degree of partial melting, 2 ) the varying degree of depletion or enrichment of the
mantle wedge from which the above partial melts are possibly generated from and 3) the
geochemical variability of the fluids released from the slab.
Figures 31 to 32 illustrate that the samples from the FCM show that fluid from the
subducting slab is affecting their geochemistry to a certain degree where as sediment
added from the subducted slab are not as important as noted by large range o f values for
Ce/Pb (1-9) and small ranges for Cs/Rb (0.01-0.08). Pb is enriched in slab fluid therefore
low Ce/Pb values represent an increase in fluid component where as Cs is enriched in
sediments relative to Rb, therefore low Cs/Rb represent a low sediment component in the
FCM.
Tectonic discrimination diagrams o f Pearce et al., (1977) which discriminate
between different geotectonic settings, intermediate samples from the FCM, Sulphur
Springs Range o f central Nevada, and the Ancestral Cascades plot within the an orogenic
setting (island arc and active continental margin) (Figure 41).
84
Figure 38: Shows the division of western North America defined by the 0.706 line (large
dot-dash line). Precambrian North America lies on the east side of the dashed line
(>0.706 Sr) and Paleozoic-Mesozoic oceanic volcanic arcs, accretionary prism complexes
and associated basinal successions accreted to western North America during Middle to
Late Jurassic arc-continent collision lies on the western side o f the dashed line (<0.706).
Figure from Streck et al., 1999.
85
■
♦
FCM
Sulphur Springs
Ancestral C asca d es
§
o
Yb
Q»
M)
ta
R>
Sr
Nd
Sm
ti
Y
Ui
Figure 39: Incompatible element diagrams for the FCM (blue), Sulphur Springs (red), and
Ancestral Cascades (green). All values are normalized to primitive mantle (Sim and
McDonough, 1988)
FCM
Ancestral Cascades
8
Basaltic Andesites Only
Figure 40: Incompatible element patterns for FCM basaltic andesites vs. Ancestral
Cascades basaltic andesites normalized to primitive mantle (Sun and McDonough, 1989).
86
On the discrimination diagrams of Pearce et al., (1984), which are based on Rb, Nb, and
Y abundances, the felsic rhyolite and dacite lavas are similar to subduction related
volcanic arc granites in central Nevada, the volcanic rocks from the East Sulphur Springs
and Eocene volcanic suites from Bingham, Utah, and along the Carlin Trend, Nevada
(Figure 42A) (Ryskamp et al., 2008). Based on Zr-Ti-Ce-P systematics, mafic rocks in
the FCM are also of a continental arc-type (Figure 42B). The most primitive basaltic
andesites of the FCM have relatively low Ti0 2 (<1.4 wt%) similar to mafic rocks found
in arcs (Figure 26) (Pearce et al., 1977). I conclude that, similar to the Sulphur Springs
Range of central Nevada, the FCM suite is related to subduction.
Having established that intraplate and mantle wedge components are not found in
the FCM, the basaltic source of the melt was likely derived from partial melting of
lithospheric mantle. However, with the large ranges o f isotopic compositions observed in
the basaltic andesites and andesites, not all intermediate rocks from the FCM can be
produced by crystal fractionation o f a calc-alkaline basaltic source from the same area so
it is clear that the FCM rock types tap a mantle source with variable isotopic composition
or are variably contaminated.
Sediments add radiogenic Sr and Pb but non-radiogenic Nd to the overlying
mantle wedge, but in different concentrations that ultimately affect the isotopic ratio of
these elements in partial melts of the wedge (Elliott et al., 1997). If sediments did not
play an important role in magmas of the FCM (low Cs/Rb on Figure 32), they are most
likely not the source of the high 207Pb/206Pb component in the FCM. Figure 32 also
suggest an enriched mantle source for the basaltic andesites where Ce/Pb ratios vary from
4 to 10 in contrast to MORB or OIB ratios of 25.
87
1 = Spreading Center Island
2 = Orogenic
FeO
3 = Ocean Ridge and Floor
4 * Ocean Island
5-Continetal
■ FCM
+ Sulphur Springs
A Ancestral Cascades
MgO
Figure 41: Ternary diagram o f FeOT, MgO and AI2O 3 discriminating between different
geotectonic settings. All samples plot within an orogenic setting (island arc and active
continental margin). Fields from Pearce et al (1977).
88
A Silicic rocks
100
Widiia Plate
400
n
Z 1
▲ D adte
♦ Rhyolite
FCM
<4> Amktite
Basaltic andesite A mafic dike*
Plagioclasc dacitc
GUotite dacite tuff
A Bkrtite porphyry
Rhyolite lava flow
Latite lava flow
Union fuff
m
I
•
10
X
wOKWaCntC
*
40
Tuacarona. Emigrant Past
Sulphur
S p rin g s
□
Carlin Trend
Ocean Ridye
a, .1, a.aa.aaai..
i
i liiiiii
10
' a » «.
100
1000
Y (ppm)
B Mafic rocks
1000
400
100
rum voaBiooift
10
• A ndesite
■ Basaltic andesite
1000
100
Ce/P20s
Figure 42: Tectonic discrimination diagrams for rocks from the FCM compared to the
volcanic suits from East Sulphur Spring, other Eocene volcanic suites from Bingham,
Utah, and along the Carlin trend, Nevada. (A) Silicic rocks (Pearce et al., 1984). (B)
Mafic rocks (Muller and Groves, 2000). Compositions of East Sulphur Springs samples,
Bingham, Carlin, Tuscarora and Emigrant Pass are from Ryskamp et al. (2008) and
sources within.
89
In this case, subduction-related metasomatism added Pb but not very much Ce to the
lithospheric mantle. Figure 43 from White (1997) plots Pb isotopic ratios in major
terrestrial reservoirs on the 208Pb/204Pb vs.
206Pb/204Pb
206Pb/204Pb
(Figure 43a) and 207Pb/204Pb vs.
(Figure 43b) plots. In both plots, FCM Pb isotopic ratios plot within the
marine sediments to upper crustal fields suggesting high radiogenic lead from crustal
contamination of crustal sediments. Sediments have concentrations o f Sr and Nd similar
to igneous rocks, but tend to be high in Pb since the sediments are derived from older
crustal rocks. Older crustal rocks tend to have higher Sr, Pb and lower Nd isotopic ratios
than the mantle. Crustal rocks also have high U/Pb ratios, which is why crustal rocks tend
to have high Pb isotope ratios. Adding U as a fluid component does not help, because
207Pb/204Pb
ratios are low in modem subducting slabs, and the characteristics of FCM
rocks is higher
A A i
Pb/
Pb than typical upper mantle. Thus the source of high
OA'7
Pb must
be old such as upper crustal rocks or possibly an older and therefore more enriched 1 Ga
Precambrian lithospheric mantle source.
In the Great Basin,
87Sr/86Sr
and Rb/Sr ratios are lower for high-silica alkali
rhyolite and Cascadia sediments than for calc-alkalic rocks (Scott et al., 1971). This
suggests a high radiogenic Sr origin of calc-alkalic rocks by partial melting of the
enriched mantle in contrast to crustal contamination of high silica crust (Yoder, 1973). If
sediments and crustal contaminants are not responsible for the higher Sr, Pb and lower
Nd isotopic values found in basaltic andesite samples from the FCM, then there must be
some source within the mantle that is responsible for the shift in isotopic values. The
importance of a lithospheric mantle source is supported by the observation that FCM
region mafic lavas commonly resemble alkaline to potassic, Miocene-Pliocene lavas from
90
Marine Sediments^
dJppar Crust
^
bower Crustal
XenoMhs
MORB
Marine Sediments/
UpperCru$t ,
iqylME
LETiON
^
Lower Crustal
Xanofiths
18
19
2t*Pb/30*Pb
^A O
1
* A J
Figure 43: Pb isotope ratios in major terrestrial reservoirs plots for A) Pb/ Pb vs.
2 ° 6 p b / 204p b a n d B ) 2 0 7 p b / 2 0 4 p b v s 2 0 6 p b / 2 0 4 p b Typical lower continental cmst and upper
continental crust are represented by lower crustal xenoliths and modem marine sediments
respectively (these somewhat underestimate the total variance in these reservoirs).
MORB and oceanic islands represent the isotopic composition of upper mantle and deep
mantle respectively. FCM symbols same as in Figure 25. Figure modified from White
(1997).
91
the Western Great Basin in terms of trace elements and radiogenic isotope ratios (Figure
36 and 37). Cousens et al. (2008) noted that the elevated 87Sr/86Sr, 207Pb/204Pb and low
143Nd/144Nd
values of the Ancestral Cascades compared to Southern Modem Cascades
mafic lavas is consistent with an enriched ancient lithospheric component beneath the
Lake Tahoe-Reno area. Higher 87Sr/86Sr compared to lavas from the Ancestral Cascades
areas may represent an older mantle due to higher ratios of Rb/Sr, or it may represent a
melt that is more modified by a lithospheric mantle component (Cousens et al., 2008).
Ancestral Cascade lavas also have mantle wedge components that are not found in the
FCM that add less radiogenic 87Sr/86Sr and more radiogenic !44Nd/143Nd to Ancestral
Cascade sources. Therefore, the addition of the mantle wedge component to the Ancestral
Cascades mantle source would have brought the 87Sr/86Sr ratio down. It is also plausible
that the lithosphere may have had a thicker metasomatized mantle or that the crustal
component may have been thicker in the central Great Basin and may be responsible for
the higher Sr isotopic values seen in the FCM Paleogene lavas (Figure 36). A thicker
crust beneath the FCM would account for the more radiogenic Sr and less elevated Nd
isotopic values compared to the Ancestral Cascades and WGB, as it may allow for
greater periods of interaction between the rising magma and may also explain why
andesitic lavas are the more dominant lava type. A thicker lithosphere may force basaltic
magmas to pond at deeper crustal levels, allowing the magmas to crystal fractionate or
mix to more evolved compositions. The increase in thickness would also allow longer
periods of interaction between the melt and crust, possibly undergoing additional crustal
contamination during magma ascent. The anomalous sample with higher than average
radiogenic Sr (10-BV-16) may be derived from a slightly more enriched overall mantle
92
source and is the only basaltic andesite found in the east portion of the FCM. Over time,
the 87Sr/86Sr in the mantle will increase due to radioactive decay of 87Rb, which may
suggest that a reason behind sample 10-BV-16 increased radiogenic Sr compared to the
rest is that the mantle source o f Horseshoe Basin may be slightly older.
Based on typical subduction signatures, high 87Sr/86Sr isotopic ratios and lower
Sr/P pmnratios, the calc-alkaline primitive sources for the FCM samples were likely
derived from partial melting o f an enriched lithospheric mantle source to generate basalt
and basaltic andesite. One way to initiate melting in the lithospheric mantle is the
addition of hydrous fluids possibly from the subducting Farallon plate, or any mantle
wedge derived magmas could cause partial melting or melt-rock reaction with
lithospheric mantle peridotite along conduit walls (Yogodzinski et al., 1996). Since FCM
isotopic geochemistry does not represent mantle wedge derived magmas typical of the
Modem South Cascades flows (Figure 36), heat and fluids necessary to partially melt the
lithospheric mantle were derived from ( 1 ) hot asthenospheric in flow impinging on the
base o f the previously metasomatized lithospheric mantle as a result of slab rollback and
(2) possible hydrous fluids released from the subducting Farallon Slab. Slab derived
fluids could either be derived from the dehydration o f the Farallon plate as hot
asthenosphere came into contact with the slab due to rollback or during flat-slab
subduction where hydrous fluids were added to the base o f the lithospheric mantle
causing metasomatism. Even though no mantle wedge evidence was found in the FCM, it
is still possible that fluids released from the subducting slab, lowered the melting
temperature of the overlying mantle wedge. The mantle wedge melt then could have
aided to melt the lithosphere but never made it past that.
93
As hot asthenosphere in-flow impinged on the base o f the previously
metasomatized lithosphere, it caused partial melting of lithospheric mantle, generating
calc-alkalic arc-like magmas that have incompatible element patterns that resemble
modem subduction-related magmas but have much higher 87Sr/86Sr and lower
143Nd/144Nd
that reflect the ancient age of the lithosphere much like isotopic ratios found
in the WGB. The relatively mafic basaltic andesitic samples that contain olivine probably
evolved from their basaltic parents mainly by fractional crystallization, as indicated by
equilibrium textures and by relatively low MgO, CaO (Figure 26) and high abundances of
SiC>2 compared with typical basalts. There is also no textural evidence for magma mixing
such as disequilibrium textures in the more mafic basaltic andesites and the presence of
olivine makes it unlikely that they formed by melting in the crust (Askren et al., 1997).
However, basaltic andesite samples 39 and 40 have greater disequilibrium textures and
lack olivine, possibly as a result of a greater fluid component or a lower degree o f partial
melting. Sample 39 and 40 also contain much lower abundances of Cr and Ni. The
basaltic parent magmas probably had the same arc-type trace element signature as the
andesitic lavas because fractionation of crystals typically found in basalts is not capable
o f depleting the magmas in Nb and Ta or producing elevated La/Nb and Ba/La ratios
(Askren et al., 1997).
7.5 Dacite and Rhyolite genesis
Dacitic and rhyolitic rocks may be formed by either the crystal fractionation o f an
andesitic melt, which is the dominant process for dacite and rhyolitic production in the
Modem Cascades (Hildreth, 2007), or by partially melting a basaltic rock. In order to
distinguish between both processes, REE plots and incompatible element patterns for
94
both dacite and rhyolite are used. Negative Eu, P, Sr and Ba anomalies (Figure 30d and
44) would represent significant crystal fractionation, as plagioclase fractionation will lead
to the Eu and to some of the Sr anomaly, where as fractionation of apatite would produce
all four anomalies. Consequently, fractional crystallization appears to have not played a
major role in forming magmas of evolved compositions represented by the lack of
anomalies in FCM lavas although dacites do have a small Eu anomaly (Figure 44).
Rhyolites from the study area were also phenocryst poor relative to the other rock types,
which according to Hildreth (2007) favours melt extraction from pluton sized crystal
mush reservoir.
Partial melting of a basaltic parent would produce a higher yet similar
incompatible element pattern and would also produce a melt with the same isotopic value
as the parental rock. Evidence which supports partial melting of crustal basaltic material
to form dacitic magma include a lack of significant Eu anomaly and similar incompatible
element pattern the dacites have compared to the basaltic andesites besides higher LREE
enrichment found in dacites (Figure 30,44). Rhyolite REE plots (Figure 44) also show a
lack of negative Eu anomaly which suggest partial melting as well; however, from a more
evolved source such as basaltic andesite in order to generate a rhyolitic melt. The
volcanic dome o f sample 10-BV-48 varies from dacite (sample A) to rhyolite (sample B)
in composition, which suggests dacitic and rhyolitic magmas were formed around the
same time.
Therefore, felsic magmas in the FCM are formed primarily by partial melting o f
middle crustal metabasaltic rocks which forms dacitic magma and by partial melting of
basaltic andesite to form rhyolitic magma (Rapp and Watson, 1995). The heat and fluids
95
required to melt the crust must have been provided by the arc-like magmas derived from
the partial melting of the lithospheric mantle. Given the lower isotopic Pb ratios for
dacites and rhyolites (Figure 33b,d) and a different isotopic trend on Figure 33a
compared to the basaltic andesites, the felsic rocks were most likely derived from a
younger possibly less enriched source such as Phanerozoic lithospheric mantle that was
previously emplaced causing partial melting of crustal rocks. This is made possibly in the
FCM since the FCM lie very close to the western margin of Precambrian lithospheric
mantle or the 0.706 Sr line. This hot magma rose buoyantly into the crust, stagnated
because of density differences in the crust, and promoted partial melting of the
continental crust to form metaluminous silicic magma. Figure 45 and 46 suggest that the
felsic FCM samples evolved through young accreted oceanic geosynclinals terranes west
o f the 0.706 Sr line suggesting metaluminous silicic magma as opposed to peraluminous
silicic magma characteristic of Precambrian continental crust to the east. FCM felsic
samples plot with Eugeocline and Sierra Nevada batholiths o f DePaolo and Farmer
(1984) on the Epsilon Nd vs. Epsilon Sr plot (Figure 45) and as metaluminous on Figure
46, both characteristic of the west side of the 0.706 Sr line. Partial melting o f middle
continental metabasaltic and metabasaltic andesite crust will produce a silicic melt that
may have impeded the ascent of more mafic magmas due to density differences resulting
in magma mixing.
96
Dacite
Rhyolite
I
Figure 44: REE diagrams for dacite and rhyolite FCM samples. All values are normalized
to primitive mantle (Sim and McDonough, 1988).
97
Primitive
Vblcanic Arcs
Eugeocline and
Sierra Nevada
O 0 - » Ms
•5 0 1 0 0
* 100—150
+ 150—100
O blot itv + muocovils
O b n t'tc thornblend*
Miogeocline
c ra to n
Figure 45: Initial eNd and eSr values o f granitic rocks in the northern Great Basin divided
by the groupings eugeocline, miogeocline and craton. Eugeocline corresponds to the
accreted area west of the 0.706 line and Miogeocline and craton correspond to areas
above the start of the Precambrian crystalline basement and the Archean Craton itself.
FCM symbols are as in Figure 25. Figure modified from DePaolo and Farmer, 1984.
<0
Peraiuminous
Metaluminous
▲ dacite
♦ rhyolite
oCM
C
M
<8
z
S'
oC
M
3
' Peratkaline
i
0.6
0.8
1.0
1.2
1.4
AI203/(CaO*Na20+K20)
Figure 46: Shands index showing mainly metaluminous compositions for FCM samples.
Fields from Shand (1943).
98
7.6 Andesites
Similarities in the geochemistry between the basaltic andesites and andesites
indicate that both are most likely derived from the same mantle source. Comparing the
trace element plots for the basaltic andesites and andesites from the FCM (Figure 30b,c),
the basaltic andesites show a trace element pattern that is very similar to the andesites;
however, with a few exceptions where the overall abundances of LILE and HFSE are
lower than the andesites. Compared to the basaltic andesites, the andesites appear to be
higher degree partial melts as they generally exhibit lower abundances of MREE and
HREE, which behave incompatibly during melting but are diluted as the percentage of
melting increases. Hornblende fractionation might also be the cause, however it would
have to be a lot of fractionation which does not appear to be the case in FCM andesites
represented by the lack of hornblende phenocrysts (< 3%).
Andesites may be produced by crustal fractionation from the basaltic andesites
melt in which case their REE patterns (Figure 47) would show a negative Eu anomaly
due to plagioclase fractionation, which is not evident, and a higher abundance of other
REE relative to basaltic andesites. Due to the lack o f significant Eu anomaly, the
andesites may be formed primarily by magma mixing between a basaltic andesite melt
with mantle-like characteristics and a felsic dacitic melt derived from partial melting of
middle cmst and/or with a rhyolitic melt derived from previously emplace mantle-derived
melt. On a Nd and Sr isotopic ratio plot (Figure 33a), the Sr isotopic values of the
andesites from the FCM show overall more radiogenic Sr values than those of the basaltic
andesites.
99
Sam ple/Prim itive M antle
* Andesites
O
Figure 47: REE patterns for Andesites normalized to Sun and McDonough (1989).
100
This leads to the proposal that the andesites may have formed by mixing a more primitive
melt that has a more depleted mantle-like isotopic signature (higher
87Sr/86Sr)
143Nd/ 144Nd,
lower
with a source that has higher Sr and lower Nd isotopic ratios such as the crust
(Cousens et al., 2008). A mixing line between these two end members would pass
through most samples in Figure 33a and produce a trace element pattern that exhibits no
large negative Eu anomaly. Mixing between a rhyolitic magma and basalt is
characteristic of samples found in the Sulphur Springs Range where quartz and olivine
phenocrysts coexist together in andesite. No such mineral phases have been found
together in the andesitic rocks of the FCM; however in terms o f disequilibrium textures,
dacitic rocks similar to andesitic rocks display sieve texture and reaction rims indicating
disequilibrium mixing or pressure-temperature change rather than equilibrium textures
characteristic o f fractional crystallization.
Textural evidence has been found to support magma mixing between mafic and
felsic magmas. Sieve texture is common on andesitic plagioclase phenocrysts (Figure 19)
of the FCM which is indicative of non-equilibrium crystallization and was likely
triggered by magma mixing within the system. This texture is also very typical of
Ancestral Cascade andesites (Cousens et al., 2008). Other disequilibrium textures include
reaction rims and extensive resorption of phases that may have come from a different
magma source. Oscillatory zoning of plagioclase present in many samples may also have
been related to the injection of a hotter, more juvenile magma into a cooling and
crystallizing chamber. The common occurrence of corroded or remelted embayments of
the crystal rims accompany many reversals supports this conclusion (Winter, 2001).
Mineral assemblages, olivine presence (in andesitic sample 10-BV-08), and trends on
101
most silica variation diagrams (Ti, Al, Fe, Mg, Ca, Na, and K) suggest it was a member
of the basaltic andesite suite. The majority of linear trends observed on two element
variation diagrams of compatible and trace element concentrations are more typical of
magma mixing as oppose to curved trends and sharp decreases more typical of fractional
crystallization (Ryskamp et al., 2008) (Figure 26). When compared to trace element
mixing plots of Ryskamp et al. (2008), die lava flows of the FCM plot in similar mixing
relationships as the samples from Sulphur Springs for incompatible Ba vs. SiC>2
suggesting they may have formed in a similar fashion; by mixing with silicic crustal
magma (Figure 48).
Therefore, it is possible that mainly hot asthenospheric in-flow on the base o f the
previously metasomatized lithospheric mantle, lowered the melting temperature o f the
lithospheric mantle, generating calc-alkaline magmas as a result of hydrous partial
melting. This hot basaltic magma rose buoyantly into the crust as it experienced rapid
decreases in pressure and fractionated olivine but no plagioclase to form basaltic
andesites. The basaltic andesites magma then intercepted and mixed with metaluminous
silicic melt that was derived from partial melting o f the crust from previously emplaced
Phanerozoic mantle-derived melt to form mainly biotite-homblende andesites. As the
overlying felsic magma either crystallized at depth or erupted, it allowed more mafic
magmas to do the same.
The crustal silicic magma may have prevented the ascent of andesitic lavas
because the density of felsic magma (2.45 g/cm 3 for dacitic magma; Whitney & Stormer,
1985) is less than that o f typical olivine or hornblende andesites (2.66 g/cm 3 at 54.6%
S i0 2 and 2.56 g/cm3 at 59.4% S i0 2; Gill, 1981) .
102
7000
Sulphur Spring Range
t
6000
Basaltic andesite and m sfe dikes
PUfioclasc decile
Biotile dame tuff
Btcxrtc porphyry
KliyoBlehmi flow
tam e lava flow
Union taff
\
Baaaitk endeeile-tliyolile mixing
% A rvlaiir mixing m od
5000
4000
\*
3000
A n d e site
Fractional cryrtxUtxatton
2000
o Bingham volcanic field
m Carlin volcanic fields
1000
0
40
45
50
55
60
65
SKh<»l%)
70
75
80
■
+
A
+
FCM
Basaltic andesite
Andesite
Dacite
Rhyolite
Figure 48: Ba vs. SiC>2 variation diagram comparing compositions of FCM samples to
those o f East Sulphur Spring, Carlin, and Bingham volcanic suites. Blue lines show the
results of mixing basaltic andesite with rhyolite, red arrows are schematic paths for
fractional crystallization of basaltic andesite, and yellow line represents the andesite
mixing trend. Figures modified from Ryskamp et al. (2008) and sources within.
103
Subsequent eruption of felsic magmas then provided unobstructed conduits for the
eruption of andesite magmas, as felsic magma chamber emptied or remaining magma
crystallized. If emptied, then felsic magma no longer remained to trap andesitic magmas.
If crystallized, then the increased density of crystallized dacitic magma (to 2.65 g/cm 3 for
a typical granodiorite) allowed hornblende andesite magma to rise and erupt (Askren et
al., 1997).
7.7 Evidence of crustal contamination
There exists evidence that the basaltic andesites and andesites may have been
affected by crustal contamination due to the presence of plagioclase crustal xenocrysts;
however no quartz xenocrysts were found suggesting little to no interaction with
Paleozoic metaluminous metasedimentary and granitic crustal rocks that are widespread
in the Great Basin (Kistler et al., 1981; Barnes et al., 2001). Assimilation of
metaluminous and peraluminous granitoids of North-central Nevada is a possible cause
o f the distinctions between FCM and WGB
87Sr/86Sr
values. Isotopically, central Nevada
granitoids have a high 87Sr/86Sr component similar to FCM lavas. Figure 49 shows that
Rb/Sr increase as magmas become more evolved and show a notable increase in
87Sr/86Sr.
Unlike the Ancestral (Tahoe-Reno) and Modem South Cascades, a rough linear mixing
relationship is found between FCM basaltic andesite and a Sierran Nevada batholiths-like
source in central Nevada. Lower degree partial melting, of central Nevada granitoids
would produce liquids potentially with even higher Rb/Sr and
87Sr/86Sr
than bulk
granitoid, since one of the first phases to melt is high Rb/Sr biotite (Kaczor et al., 1988;
Knesel and Davidson, 1996).
104
0.710
0.709-
Enriched mantle
source
0.708CO
<o 0.707-
★ +as i«'fta'j0
%
co
h- 0.70600
.
A Tahoo-Reoo
0.705-
O South Cascades
i f Sidrran Granites
0.704-j
Fractionation
0.7034
Rb/Sr
Figure 49: Initial Sr isotope ratios vs. Rb/Sr plot for FCM samples plotted along with
Sierran Granites, Ancestral Cascades Tahoe-Reno area and Modem South Cascades.
FCM symbols are the same as in Figure 25. Sierran Granites, Tahoe-Reno and South
Cascades geochemistry from Cousens et al. (2008).
105
However, lower 87Sr/86Sr and Rb/Sr than granitoids combined with no negative Eu
anomaly from the melting of feldspar suggest no assimilation of central Nevada
granitoids by FCM magmas (Figure 49; Cousens et al., 2008).
The 87Sr/86Sr from the FCM are more radiogenic than the Ancestral Cascades and
overlap with the most radiogenic lavas of the Western Great Basin suite (Omerod, 1991).
The more evolved rocks of the FCM approach crustal 87Sr/86Sr ratios, suggesting Sr
contamination from crustal sources (Figure 36). SiC>2 concentrations and
87Sr/86Sr
are
useful geochemical tools in helping to identify continental crust as a possible contaminant
of ascending magmas. On the Si0 2 vs.
87Sr/86Sr
plot (Figure 50), there is a notable
increase in radiogenic Sr with increasing Si0 2 and a notable decrease in radiogenic Nd
with increasing Si0 2 (Figure 33e) which may suggest contamination. However, the
positive correlation of Figure 50 is more likely to represent mixing o f rhyolitic magma
with basaltic andesite in order to generate dacites and rhyolites. The positive correlation
that Pb and Th have with Si0 2 may indicate that FCM samples experienced some Pb and
Th contamination in contrast to other incompatible elements such as Ba or Zr which vaiy
irregularly (Figure 28).
106
s
©
to
K
o
s-
o
N.
o
andesite
dacite
Crust
o
8
©<0 ^O
K
o
§
L mantie
o
r^
55
60
65
70
Si02
Figure 50: S1O2 vs. radiogenic
are as in Figure 25.
87Sr/86Sr plot.
L mantle = Lithospheric mantle. Symbols
107
7.8 Mantle beneath the FCM and tectonic model
Tectonic setting discrimination diagrams clearly show that FCM lava flows are
calc-alkaline on Figure 35 from Wood (1980) which are widespread in the central Great
Basin (Best et al., 1989). Figures43A (felsic) and B (mafic) show that samples are related
to a continental volcanic arc suite presumably related to the rollback o f the subducting
Farallon plate under North America.
Late Cretaceous subduction related magmatism ended near Lake Tahoe around 90
Ma as the magmatic activity o f the Sierra Nevada batholith shifted eastward, most likely
as a result of shallowing o f the dip of the subducting Farallon slab (Dickinson and
Snyder, 1978; Lipman, 1992). During flat slab subduction, metasomatism changed the
base of the lithosphere and added components enriched in LILE, LREE, Zr and depleted
in Nb and Ti, which are characteristics expected in components derived from the
subducting slab (Rivalenti et al., 2007). Given the Ar-Ar ages determined from the
basaltic andesite to rhyolites of the FCM, they span approximately 1 m.y. from 33.3 to
34.3 Ma during the Paleogene time period. At this time the Great Basin experienced a
significant change in magmatism when the Farallon plate detached from the lithosphere
in a southwestward sweeping fashion allowing hot asthenospheric mantle to flow
between the subducting slab and the lithospheric mantle (Best and Christiansen, 1991).
This created a continental magmatic arc that also swept to the SW producing magma
compositions ranging from basaltic andesite to rhyolite (or their intrusive equivalents)
(Severinghaus and Atwater, 1990). Middle Paleogene magmatism may have resulted
from ( 1 ) decompression melting of hot mantle as a result o f small scale convection above
the subducting Farallon plate, (2) dehydration of the Farallon plate, which caused
108
hydrous melting o f the mantle wedge, and (3) heating of metasomatized lithospheric
mantle by hot asthenospheric mantle or by wedge derived magma. These mafic, mantlederived magmas ascended and powered crustal magma systems in which more felsic
magmas evolved by partial melting continental crust and magma mixing (Ressel and
Henry, 2006).
From studying the most primitive basaltic andesites and andesites, it is clear that
there is a distinctive geochemical domain in the mantle. These distinctive signatures are
identifiable when looking at the incompatible element plots and radiogenic isotopes
(Figure 30, 34a,b,d). There are mainly two geochemical components present in the
mantle beneath the FCM. These are an enriched ancient Precambrian lithospheric mantle
source metasomatized by previous subduction episodes to generate the basaltic andesites
combined with mixing of crustal magma which has parental melts derived from a less
enriched younger Phanerozoic mantle lithosphere.
FCM Eocene-Oligocene lava flows form a dominantly high K, calc-alkaline suite
with low Mg #’s similar to those found in subduction settings worldwide. The volcanic
and subvolcanic rocks range from olivine-bearing, high-MgO basaltic andesite to high
silica rhyolite. The intermediate to silicic rocks have subduction related trace element
signatures with Nb-Ti depletions and enrichment o f Pb, also similar to those formed at
convergent margins, and have highly radiogenic87Sr/86Sr.
This leads to the assumption that melts are being stored as batch melts in the
middle crust after partially melted from enriched ancient lithospheric mantle and
109
undergoing fractional crystallization processes from basalt to basaltic andesite given the
presence of olivine, Cr-spinel and equilibrium textures in basaltic andesites.
These characteristics suggest the magma system was rooted in a subduction zone
setting that formed as the Farallon plate steepened during the Paleogene where heating
and partial melting of lithospheric mantle by hot asthenospheric mantle and possibly
wedge derived magmas was the likely cause of middle Paleogene magmatism in the FCM
(Figure 51). As hot asthenosphere came in contact with the base of the previously thick
metasomatized lithospheric mantle, arc-like magmas with elevated 87Sr/86Sr, low
143Nd/ 144Nd,
and high 207Pb/204Pb were generated as a result o f hydrous partial melting of
Precambrian lithospheric mantle. Small proportions of mantle wedge derived magma
could have been generated from slab fluids caused by the dehydration of the subducting
Farallon slab however; FCM geochemistry does not show evidence to support this. It is
possible that this mantle wedge derived magma helped to partially melt the lithospheric
mantle but did not make it past the base of the lithosphere. The more mafic sub-alkaline
samples 16 and 42 are most likely derived from higher degrees of partial melts given
their more subalkaline nature where as samples 39 and 40 are alkaline.
This hot magma ascended buoyantly into the crust as it experienced rapid
decreases in pressure, intercepted and mixed with previously emplaced rhyolitic magma
derived by partial melting of crustal basaltic andesites to form mainly biotite-homblende
dacites and andesites. Petrographic evidence, linear trends on variation diagrams,
homogeneous incompatible element plots and overlapping isotope plots support this
conclusion. In addition, evidence of magma mixing involving alkaline olivine bearing
magmas is widespread in the central Great Basin. During the Eocene, mafic alkaline
110
magmas intercepted and mixed with evolved calc-alkaline magmas in shallow
subvolcanic settings (Ryskamp et al., 2008). Structural boundaries including Proterozoic
age basement-penetrating faults, Paleozoic and Mesozoic thrust faults, and magma­
generated fractures, guided magma emplacement routes through thick central Great Basin
continental crust and controlled levels of stagnation. As the more evolved magmas
crystallized at depth or erupted, the increased density of crystallized dacitic magma or
eruption of overlying felsic magma allowed andesite and basaltic andesite magma to rise
and erupt. As a result of crustal trapping, only a small fraction of the mafic magma was
able to erupt (Figure 51).
I ll
A. 120 to 45 Ma F1* Stab SUbductioo
W
California
Nevada
Utah
E
OT RMT
Figure 51: Plate tectonic reconstruction of the western United States modified from
Ryskamp et al. (2008) during the Paleogene Period. Gt—Golconda thrust; RMT—
Roberts Mountains thrust. (A) A period of low-angle subduction during the Cretaceous
and early Paleogene metasomatises the lithospheric mantle, thickens the crust and magma
generation is stalled. (B) The subducting oceanic lithosphere rolled back inducing
asthenosphere counterflow. As it heated, the slab dehydrated and initiated the production
o f hydrous, oxidized fluids in the overlying mantle wedge over a broad area (green). Hot
asthenospheric mantle impinges on the base o f the enriched lithospheric mantle near the
edge of the Precambrian continental crust inducing partial melting of lithospheric mantle.
The melts ascend and undergo fractional crystallization to form basaltic andesite from
basalt. Basaltic andesite magma intercepts and mixes with rhyolitic magma derived from
partial melting o f crustal metabasaltic andesites to form dacites and andesites. Structural
boundaries including faults and magma generated fractures guided magma emplacement
routes and controlled levels of stagnation.
112
Conclusions
Paleogene volcanic rocks collected from the Fish Creek Mountains and
surrounding Tobin and Shoshone Range area in north-central Nevada, are calc-alkaline
basaltic andesites through to rhyolites and range in age from 33.3-34.3 Ma. Volcanic
edifices include lava flows, domes and volcaniclastic flows which typically include
phenocrysts of plagioclase, pyroxene, potassium feldspar, olivine (intermediate samples)
and quartz (felsic samples), with less common phenocrysts o f hornblende, biotite, and an
unidentified opaque mineral, probably a Fe-Ti oxide.
FCM Paleogene magmatism is related to the southwestward sweep of magmatism
due to the increased angle of subduction of the ancient oceanic Farallon plate under North
America. There are mainly two geochemical components present in the mantle beneath
the FCM, which are an enriched ancient Precambrian lithospheric mantle source and a
younger less enriched Phanerozoic lithospheric mantle responsible for the partial melts o f
rhyolitic crustal magma.
Major and trace element geochemistry suggest that FCM volcanic compositions
are generally consistent with a subduction zone origin including low Mg #’s, high K 2O,
negative Nb and Ti anomalies and high concentrations of LILE to HFSE. The enrichment
of mobile elements (LILE) to depletions of immobile elements (HFSE) reflect melts of
the mantle sources that are metasomatized by hydrous fluids from the subducting slab.
Middle to heavy REE patterns suggest the melt was generated at depths within the spinel
lherzolite field (<75 km depth).
113
Crustal contamination processes in the FCM magmas did not play an important
role and are not the source of high
87Sr/86Sr, 207Pb/206Pb
and low l43N d/144Nd in the FCM.
Rather, low Sr/Ppmn values and high initial 87Sr/86Sr ratios suggest the FCM suite to be
derived from enriched components in a 1 Ga Precambrian lithospheric mantle which was
metasomatized by earlier subduction episodes, similar to the source o f the WGB. If an
enriched ancient lithospheric mantle source existed beneath the FCM region as in the
WGB, magmas derived from this source would have incompatible element patterns that
resemble modem subduction-related magmas but have higher 87Sr/86Sr and lower
143Nd/,44Nd
that reflect the ancient age o f the lithosphere.
Samples from the FCM exhibit values of 87Sr/86Sr that are slightly more
radiogenic compared to the Ancestral Cascades (Cousens et al., (2008) and most of the
Western Great Basin (Ormerod et al., 1991). This is because there are no mantle wedge
components in the FCM and possibly an older lithospheric mantle. It is also plausible that
the lithospheric component may have been thicker in the central Great Basin as a result of
flat-slab subduction and may be responsible for the higher Sr isotopic values seen in the
FCM Paleogene lavas. A thicker lithosphere beneath the FCM would account for the
more radiogenic Sr and less elevated Nd isotopic values compared to the Ancestral
Cascades and WGB, as it may allow for greater periods of interaction between the rising
magma and may also explain why andesitic lavas are the more dominant lava type.
Therefore, partial melting of an enriched Precambrian lithospheric source with no mantle
wedge component could have occurred in order to get the elevated 87Sr/86Sr, low
143Nd/ 144Nd,
and high 207Pb/204Pb values seen in the FCM basaltic andesites through to
rhyolites.
114
Fractional crystallization appears to have played a minor role due to the lack of
significant Eu anomaly in FCM REE plots and suggests it was not the dominant process
in forming magmas o f evolved compositions. Evidence which supports partial melting of
crustal basaltic andesites to form rhyolitic magma followed by subsequent magma mixing
between rhyolite and basaltic andesite include the lack of significant Eu anomaly that
would result from plagioclase fractionation, the homogeneous incompatible element
patterns from basaltic andesite to rhyolite and the overlap between isotopic ratios. The
heat and fluids required to melt the crust must have been provided by the arc-like
magmas derived from the partial melting of the Precambrian lithospheric mantle. This hot
magma rose buoyantly into the crust as it fractionated from basalt to basaltic andesite,
intercepted and mixed with a rhyolitic magma chamber derived from previously
emplaced Phanerozoic mantle-derived melts to form biotite-homblende andesites and
dacites. The linear trends observed on two element variation diagrams o f compatible and
trace element concentrations are also more typical o f magma mixing as oppose to curved
trends and sharp decreases more typical of fractional crystallization. The melt then rose
into shallower and eventually erupted as cones (10-BV-48) or flows.
115
References
Armstrong, R.L., Taubeneck, W.H. and Hales, P.O., 1977. Rb/Sr and K/Ar chronometry
of the Mesozoic granite rocks and their Sr isotope compositions, Oregon,
Washington, and Idaho. Geological Society o f America Bulletin, 8 8 : 397-411.
Askren, D.R., Roden, M.F. and Whitney, J.A., 1997. Petrogenesis of Paleogene Andesite
Lava flows Interlayered with Large-Volume Felsic Ash-Flow Tuffs of the
Western USA. Journal of Petrology, 38: 1021-1046.
Bacon, C.R., Bruggman, P.E., Christiansen, R.L. and Clynne, M.A., 1997. Primitive
Magmas at five Cascades volcanic field: Melts from hot, heterogeneous sub-arc
mantle. Canadian Mineralogist, 35: 397-423.
Barnes, C.G., Burton, B.R., Burling, T.C., Wright, J.E. and Karlsson, H.R., 2001.
Petrology and geochemistry of the Late Eocene Harrison Pass pluton, Ruby
Mountains Core Complex, northeastern Nevada. Journal of Petrology, 42: 901929.
Bartley, J.M., Axen, G.J., Taylor, W.J. and Fryxell, J.E., 1988. Cenozoic tectonics o f a
transect through eastern Nevada near 38 N latitude. In: D.L Weide and M.L Faber
(Editors), Geological journeys in the southern Basin and Range Cordilleran
Section Field Trip Guidebook. Geological Society o f America, Boulder,
Colorado, pp. 1-20.
Best, M.G. and Christiansen, E.H., 1991. Limited Extension During Peak Paleogene
Volcanism, Great Basin o f Nevada and Utah. Journal o f Geophysical Research,
96: 13,509-13,528.
Best, M.G., Christiansen, E.H. and Blank Jr, R.H., 1989. Oligocene caldera complex
and calc-alkaline tuffs and lavas of the Indian Peak volcanic field, Nevada and
Utah. Geological Society of America Bulletin, 101: 1076-1090.
Brenan, J.M., Shaw, H.F., Ryerson, F.J. and Phinney, D.L., 1995. Mineral-aqueous fluid
partitioning of trace elements at 900 C and 2.0 GPa: Constraints on the trace
element geochemistry o f mantle and deep crustal fluids. Geochemical
Cosmochimca Acta, 59: 3331-3350.
Borg, L.E., Blichert-Toft, J. and Clynne, M.A., 2002. Ancient and modem subduction
zone contributions to mantle sources o f lavas from the Lassen region o f California
inferred from Lu-Hf isotopic systematic. Journal of Petrology, 43: 705-723.
116
Borg, L.E., Clynne, M.A. and Bullen, T.E., 1997. The variable role o f slab-derived fluids
in the generation o f a suite of primitive calc-alkaline lavas from the southernmost
Cascades, California. Canadian Mineralogist, 35: 425-452.
Christiansen, E.H., Cenozoic Magmatism in the Great Basin: A view from the East
(abstract). In: Geological Society o f Nevada Symposium; 2010 May 14-22; Reno,
NV. Provo (UT): Department of Geological Sciences; 2010. p 15. Abstract 1501.
Clark, C.A., 2011. A Geochemical Exploration of the Sagehen Volcanic Centre, TruckeeTahoe Region, California, U.S.A. Department of Earth Science. Ottawa, Calreton
University: 144
Cline, M.A., 1990. Stratigraphic, lithologic and major element chemical constraints on
magmatic evolution at Lassen Volcanic Centre, California. Journal of
Geophysical Research, 95: 19651-19669.
Cline, J.S., Hofstra, A.H., Muntean, J.L., Tosdal, R.M. and Hickey, K.A., 2005. Carlintype gold deposits in Nevada: Critical geologic characteristics and viable models.
Economic Geology, 100th Anniversary Volume: 451-484.
Cousens, B.L., 1996. Magmatic evolution of Quaternary mafic magmas at Long Valley
Caldera and the Devils Postpile, California: Effects of crustal contamination on
lithospheric mantle-derived magmas. Journal of Geophysical Research 101,
27673-27689.
Cousens, B., Prytulak, J., Henry, C., Alcazar, A. and Brownrigg, T., 2008. Geology,
geochronology, and geochemistry o f the Miocene-Pliocene Ancestral Cascades
arc, northern Sierra Nevada, California and Nevada: The roles o f the upper
mantle, subduction slab, and the Sierra Nevada lithosphere. Geosphere, 4(5): 829853.
Davies, G.F., 1999. Dynamic Earth: Plates, Plumes and Mantle Convection. Cambridge
University Press, New York. 472p.
DePaolo, D.J. and Farmer, G.L., 1984. Isotopic data bearing on the origin of Mesozoic
and Paleogene granitic rocks in the Western United States. Philosophical
Transactions of the Royal Society A, 310: 743-753.
Dickinson, W.R., 2006. Geotectonic evolution of the Great Basin. Geosphere, 2: 353368.
Dickinson, W.R., 1979. Mesozoic forarc basin in central Oregon. Geology, 7: 166-170.
117
Dickinson, W.R. and Snyder, W.S., 1978. Plate tectonics of the Laramide orogeny. In: V.
Matthews, (Editor), Laramide folding associated with basement block faulting in
the western United States. Geological Society of America Memoir, 151: pp. 355—
366.
Ellam, R.M. and Hawkesworth, C.J., 1988. Is average continental crust generated at
subduction zones?. Geology 16: 314-317.
Elliott, T., Plank, A., Zindler, A., White, W. and Bourdon, B., 1997. Element transport
from slab to volcanic front at the Marianas arc. Journal of Geophysical Research,
102(B7): 14991-15019.
Emmons, D.L. and Eng, T.L., 1995. Geologic Map of the McCoy Mining District, Lander
County, Nevada. Nevada Bureau of Mines and Geology, Map 103, scale 1:
12, 000.
Ewart, A., 1979. A review of the mineralogy and chemistry of Paleogene-Recent dacitic,
latitic, rhyolitic, and related salic volcanic rocks. In: Barker, F., (Editor),
Trondhjemites, dacites, and related rocks. Elsevier, Amsterdam, pp. 13-121.
Farmer, G.L., Glazner, A.F. and Manley, C.R., 2002. Did lithospheric delamination
trigger late Cenozoic potassic volcanism in the southern Sierra Nevada,
California?. Bulletin o f the Geological Society of America, 114: 754-768.
Faure, G., 1986. Principles of Isotope Geochemistry (second edition). Wiley & Sons,
New York.
Fiero, B., 1986. Geology o f the Great Basin. University of Nevada Press, Reno and Las
Vegas.
Gill, J.B., 1981. Orogenic Andesites and Plate Tectonics. Springer-Verlag, Berlin.
Gonsior, Z.J., 2006. The Timing and Evolution of Cenozoic Extensional Normal Faulting
in the Southern Tobin Range, Pershing County, Nevada. Department o f Earth
Science. Corvallis, Oregon State University, pp. 1-64.
Hamilton, W.B., 1994. Subduction systems and magmatism. Geological Society of
London, Special Publications, 81: 3-28.
Hawkesworth, C.J., Kempton, P.D., Rogers, N.W., Ellam, R.M. and van Calsteren, P.W.,
1990. Continental mantle lithosphere and shallow level enrichment processes in
the Earth’s mantle. Earth and Planetary Science Letters 96:256-268.
Hildreth, W., 2007. Quaternary magmatism in the Cascades-Geological perspectives.
U.S. Geological Survey Professional Paper, 1744: p. 125.
118
Hughes, C.J., 1973. Spilites, keratophyres, and the igneous spectrum. Geological
Magazine, 6: 513-527.
Humphreys, E.D., 1995. Post-Laramide removal of the Farallon slab, western United
States. Geology 23: 987-990
Humphreys, E., Erslev, E., Atwater, T., Hessler, E., Dueker, K. and Farmer, G. L., 2003.
How laramide-age hydration of North American Lithosphere by the Farallon Slab
controlled subsequent activity in the Western United States. International Geology
Review 45: 575-595.
Ida, Y., 1983. Convection in the mantle wedge above the slab and tectonic processes in
sibduction zones. Journal of Geophysical Research, 88(B9): 7449-7456.
Irvine, T.N. and Barager, W.R.A., 1971. A guide to the chemical classification of the
common volcanic rocks. Canadian Journal of Earth Sciences 8: 523-548.
Iwansson, J. K. and Landstrom, O., 2000. Contamination of rock samples by laboratory
grinding mills. Journal of Radioanalytical and Nuclear Chemistry. 244: 609-614.
Jarrard, R.D., 1986. Relations among subduction parameters. Reviews o f Geophysics.
24(2): 217-284
John, D.A., 2001. Miocene and Early Pliocene Epithermal Gold-Silver Deposits in the
Northern Great Basin, Western United States: Characteristics, Distribution, and
Relationships to Magmatism. Economic Geology, 96: 1827-1853.
John, D.A., Henry, C.D. and Colgan, J.P., 2008. Magmatic and tectonic evolution o f the
Caetano caldera north-central Nevada: A tilted mid-Paleogene eruptive center and
source o f the Caetano Tuff. Geosphere, 4: 75-106
Kaczor, S.M., Hanson, G.N. and Peterman, Z.E., 1988. Disequilibrium melting of granite
at the contact with a basic plug: A geochemical and petrographic study. Journal o f
Geology, 96: 61-78.
Kepper, J.C., 1989. Regional Structural Blocks, Volcanism, and Precious Metal Deposits
in the South-Central Nevada Structural Zone. International Geoscience and
Remote Sensing Symposium, 3: 1421-1423.
Kistler, R.W., Ghent, E.D. and O’Neill, J.R., 1981. Petrogenesis of garnet two-mica
granites in the Ruby Mountains, Nevada. Journal o f Geophysical Research, 86:
10,591-10,606.
Knesel, K.M. and Davidson, J.P., 1996. Isotopic disequilibrium during melting o f granite
and implications for crustal contamination o f magmas. Geology, 24: 243-246.
119
La Tour, T.E., 1989. Analyses of rocks using X-Ray fluorescence spectrometry. Rigaku
Journal, 6: 3-9
Lipman, P.W., 1992. Magmatism in the Cordilleran United States; progress and
problems. In: B.C. Burchfiel, P.C. Lipman, and M.L. Zoback (Editors), The
Cordilleran orogen: Conterminous U.S.: Boulder, Colorado. Geological Society of
America; Geology o f North America, G-3: pp. 481—514.
LeBas, M.J., LeMaitre, R.W., Streckeisen, A. and Zanettin, B., 1986. A Chemical
Classification of Volcanic Rocks Based on the Total Alkali-Silica Diagram.
Journal o f Petrology, 27: 745-750.
Manduca, C.A., Silver, L.T. and Taylor, H.P., 1992. 87 Sr/ 86 Sr and 18 O/ 16 O isotopic
systematic and geochemistry of granitoid plutons across steeply-dipping boundary
between contrasting lithospheric blocks in western Idaho. Contribution to
Mineralogy and Petrology, 109: 355-372.
McKee, E.H. and Silberman, M.L., 1970. Geochronology o f Paleogene Igneous Rocks in
Central Nevada. Geological Society of America Bulletin, 81: 2317-2328.
McKee, E.H., Silberman, M.L., Marvin, R.E. and Obradovich, J.D., 1971. A summary of
radiometric ages o f Paleogene volcanic rocks in Nevada and eastern California —
Part 1, central Nevada. Isochron/West, no 2: 21-42.
Ormerod, D.S., Hawkesworth, C.J., Rogers, N.W., Leeman, W.P. and Menzies, M.A.,
1988. Tectonic and magmatic transitions in the Western Great Basin, USA.
Nature, 333: 349-352.
Ormerod, D.S., Rogers, N.W. and Hawkesworth, C.J., 1991. Melting in the lithospheric
mantle: Inverse modeling of alkali-olivine bsalts from the Big Pine Volcanic
Field, California. Contributions to Mineralogy and Petrology, 108: 305-317.
Pearce, J.A., 1983. Role o f the Sub-continental Lithosphere in Magma Genesis at Active
Continental Margins. Continental Basalts and Mantle Xenoliths. C.J.
Hawkesworth and M.J. Norry. Cambridge, MA, Birkhauser Boston Inc.
Pearce, J.A., Harris, N.B. and Tindle, A.G., 1984. Trace element discrimination diagrams
for the tectonic interpretation of granitic rocks. Journal of Petrology, 25: 956-983.
Pearce, T.H., Gorman, B.E. and Birkett, T.C., 1977. The relationship between major
element chemistry and tectonic environment of basic and intermediate volcanic
rock. Earth and Planetary Science Letters, 36: 121-132.
Peccerillo, A. and Taylor, S. R., 1976. Geochemistry of Eocene calc-alkaline volcanic
rocks from the Kastamonu area, Northern Turkey. Contributions to Mineralogy
and Petrology, 58: 63-81.
120
Plank, T. and Langmuir, C., 1993. The chemical composition of subducting sediment and
its consequence for the crust and the mantle. Chem. Geol, 145: 325-394.
Prytulak, J., Vervoort, J.D., Plank, T. and Yu, C., 2006. Astoria Fan sediments, DSDP
site 174, Cascadia Basin: Hf-Nd-Pb constraints on provenance and outburst
flooding. Chemical Geology 233, 276-292
Rapp, R. P. and Watson, E. B., 1995. Dehydration melting of metabasalt at 8-32 kbar:
implications for continental growth and crust-mantle recycling. Journal of
Petrology, 36: 891-931.
Rea, D.K. and Ruff, L.J., 1996. Composition and mass flux o f sediment entering the
worlds subduction zones: Implications for global sediment budgets, great
earthquakes and volcanism. Earth and Planetary Science Letters, 140: 1-12.
Ressel, M.W. and Henry, C.D., 2006. Igneous geology o f the Carlin Trend, Nevada:
Development of the Eocene plutonic complex and significance for Carlin-type
gold deposits. Economic Geology, 101: 347-383.
Rivalenti, G., Mazzucchelli, M., Zanetti, A., Vannucci, R., Bollinger C., Hemond, C. and
Bertotto, G.W., 2007. Xenoliths from Cerro de los Chenques (Patagonia): An
example of slab-related metasomatism in the backarc lithospheric mantle. Lithos,
99: 45-67.
Rutherford, M.J., Devine, J.D. and Barclay, J., 1998. Changing magma conditions and
ascent rates during the Soufriere Hills eruption on Montserrat. GSA Today, 8:1-7.
Ryskamp, E.B., Abbott, J.T., Christiansen, E.H., Keith, J.D., Vervoort, J.D. and Tingey,
D.G., 2008. Age and petrogenesis of volcanic and intrusive rocks in the Sulphur
Spring Range, central Nevada: Comparisons with ore-associated Eocene magma
systems in the Great Basin. Geosphere, 4: 496-519
Schmidt, M.E., Grander, A.L. and Rowe, M.C., 2008. Segmentation o f the Cascades Arc
as indicated by Sr and Nd isotopic variation among diverse primitive basalts.
Earth and Planetary Science Letters, 266: 166-181.
Scott, R. B., Nesbitt, R W., Dasch, E J. and Armstrong, R L., 1971. A strontium isotope
evolution model for Cenozoic magma genesis, eastern Great Basin, U.S.A.
Bulletin of Volcanology, 35: 1-26
Severinghaus, J., and Atwater, T. M., 1990. Cenozoic geometry and thermal state of the
subducting slabs beneath North America. In: B.P. Wernicke (Editor), Basin and
Range extensional tectonics near the latitude o f Las Vegas, Nevada. Geological
Society of America Memoir, 176: pp. 1-22.
121
Shand S.J., 1943. Eruptive Rocks; Their Genesis, Composition, Classification, and their
Relation to Ore Deposits, with a chapter on Meteorites (revised second edition).
Hafner Publishing Company, New York.
Stevens, C.H., Stone, P., Dunne, G.C., Greene, D.C., Walker, J.D. and Swanson, B.J.,
1997. Paleozoic and Mesozoic evolution o f east-central California. International
Geology Review, 39: 788-829.
Stewart, J.H. and Carlson, J.E. 1976. Geologic Map of North-Central Nevada, Nevada
Bureau o f Mines and Geology; U.S. Geological Survey, Map 50, scale 1:250,000.
Stoffers, A.J., 2010. Geochemical Transition from Miocene-Pliocene to Quaternary Arc
Volcanism in the Sierra Nevada, Northern California. Department o f Earth
Sciences. Ottawa, Carleton University: 128.
Streck, M.J. and Grunder, A.L., 1999. Enrichment of basalt and mixing o f dacite in the
rootzone of a large rhyolite chamber: inclusions and pumices from Rattlesnake
Tuff, Oregon. Contributions to Mineralogy and Petrology, 136: 193-212.
Sun, S. S. and McDonough, W.F., 1989. Chemical and isotopic systematic o f oceanic
basalts: Implications for mantle composition and processes. In: A.D. Saunders
and M.J. Norry (Editors), Magmatism in the Ocean Basins, Special Publication.
Geological Society of London, London, pp. 313-345.
Todt, W., Cliff, R.A., Hanser, A. and Hofmann, A.W., 1984. 202Pb -205Pb spike for Pb
isotopic analysis. Terra Cognita 4, 209.
Winchester, J.A. and Floyd, P. A., 1977. Geochemical discrimination o f different magma
series and their differentiation products using immobile elements. Chemical
Geology, 20: 325-343.
Varve, S., 2013. Stratigraphy, Geochemistry, and Origin of the Fish Creek Mountains
Tuff, Battle Mountain Area, North Central Nevada, USA. Department of Earth
Sciences. Ottawa, Carleton University.
Wetmore, S., 2011. Petrography, Geochemistry and Isotpic analyses o f the Buffalo
Vallyey Cinder Cones in the Fish Creek Mountains, North-central Nevada.
Department of Earth Sciences. Ottawa, Carleton University: 72.
White, M.W., 2011. Geochemistry. John-Hopkins University Press, Cornell University.
Whitney, J.A. and Stormer, J.C., 1985, Mineralogy, petrology, and magmatic conditions
from the Fish Canyon Tuff, central San Juan volcanic field, Colorado. Journal of
Petrology, 26: 726-760.
122
Winter, J.D., 2001. An Introduction to Igneous and Metamorphic Petrology. Prentice
Hall, London. 697p.
Wood, D.A., 1980. The application of a Th-Hf-Ta diagram to problems of
tectonomagmatic classification and to establishing the nature of crustal
contamination of basaltic lavas of the British Paleogene Volcanic Province. Earth
and Planetary Science Letters, 50: 11-30.
Yoder, H.S., 1973. Contemporaneous basaltic and rhyohtic magmas. American
Mineralogist, 58: 153-171.
Yogodzinski, G.M., Naumann, T.R., Smith, E.I. and Bradshaw, T.K., 1996. Evolution of
a mafic volcanic field in the central Great Basin, south central Nevada. Journal of
Geophysical Research. 101: 17425-17446.
Zhang, S.H., Zhao, Y., Song, B., Hu, J.M., Lui, S.W., Yang, Y.H., Chen, F.K., Lui, X.M.
and Lui, J., 2009. Contrasting Late Carboniferous and Late Permian-Middle
Triassic intrusive suites from the northern margin of the North China craton:
Geochronology, petrogenesis, and tectonic implications. Geological Society of
America Bulletin, 121: 181-200.
123
Appendix A: Rock descriptions and TAS name with sample num ber and location.
UTM: NAD 27, Zone 11
Sample #
Hand Specimen Description
Easting
N orthing
10-BV-05
Andesite. Taken from part o f the float
463137E
4458312N
465592E
4459389N
465435E
4452055N
below an outcrop. Dark grey-black
matrix. Dark green altered mafic
minerals 5-8%, feldspar laths l-2mm
crystals 10-15% and olivine (2%).
Approximately 1.5km from Jersey
summit
10-BV-06
Trachydacite. Taken from a weathered
outcrop o f plag-phyricaa or flow top
lava. Altered feldspar laths in a flow
banded reddish-brown matrix, 5-10%
plagioclase laths 2-3 mm, iron stained
fractures. Located south of Buffalo
Valley cones.
10-BV-07
Trachyandesite. From part o f the float
beneath an outcrop. Dark grey glassy
matrix, plag-phyric with 20%
plagioclase laths 1mm with minor
124
olivine (1-2%) and pyroxene. Brownred weathered surface, massive and
non-vesicular. Dated at 33.3 +/- 0.05
Ma. Located in Cow Canyon
10-BV-08
Trachyandesite. Dark grey plagioclase
465256E
4451873N
465948E
4451890N
465950E
4451892N
rich matrix with 20-25% plagioclase
laths of 1-2 mm, minor pyroxene and
rare large hornblende 4-5 mm.
Located in Cow Canyon at a higher
elevation than 07.
10-BV-09
Trachyandesite taken from a jointed
outcrop. Fine grained dark grey to
black matrix with pyroxene
phenociysts. Outcrop is massive and
non-vesicular with flat plagioclase
laths showing trachytic texture.
Located in Cow Canyon.
10-BV-10
Trachyandesite exposed in an old river
bed. Outcrop is jointed and massive
with a grey to dark grey matrix and an
oxidized weathered surface. Contains
medium grained pyroxenes (3-5mm)
125
and some olivine with elongated
plagioclase (1mm). Located in Cow
Canyon.
10-BV-l 1
Trachydacite taken from a jointed
462113E
4446107N
463056E
4446260N
463056E
4446260N
outcrop. Light grey massive and non
vesicular matrix with green altered
mafic minerals (25%), blocky laths of
plagioclase (35%) up to 15mm and
chloritized biotite (1-2%) up to 2mm.
Located in Jersey Canyon.
10-BV-12A
Rhyolite. Taken from a volcanic
breccia bedded outcrop with some
planar features. Fine grained dark grey
matrix with elongated euhedral
plagioclase laths (20%) up to l-2mm.
Some fine grained plagioclase, pyrite
grains (1%) and pyroxene (8%).
Located in upper Jersey Canyon.
10-BV-12B
Trachyandesite. Taken from a
volcanic breccia outcrop. Lighter grey
partially weathered sample of matrix,
plagioclase feldspar rich (40%), quartz
126
and accessory red garnet. Also located
in the upper Jersey Canyon.
10-BV-l 6
Basaltic Andesite taken from an
480523E
4459155N
488461E
4457888N
angular outcrop with some float. Black
fine grained matrix with pyroxene
clusters (glassy dark green 3mm)
olivine and blocky plagioclase (1%)
up to 1mm. This flow seems to be
restricted to this part, flanked by both
sides by glassy flows o f rhyolitic 10BV-15.
10-BV-29
Trachyandesite taken from a bedded
and sheared outcrop. Beds strike east
and dip 10° N. Porphyritic andesite
with elongated euhedral plagioclase up
to 5 mm (25%) and 6-7 mm
hornblende (1-5%) in a dark black
matrix. Weathered outcrop with
altered mafic minerals and some iron
stained fractures. Located at the south
end of Horseshoe Basin.
127
10-BV-30
Trachyandesite taken from the flows
488050E
4457894N
486490E
4459327N
486292E
4459257N
overlying those of sample 10-BV-29.
Light grey-green matrix with 25%
plagioclase, hornblende, 15%
pyroxene. Outcrop beds are 2-5 cm
thick. Located at the southern end of
Horseshoe Basin.
10-BV-31
Trachyandesite taken from a massive
outcrop with no apparent bedding.
Glassy medium grey matrix, round to
elongated vesicles <lmm (5%), and
euhedral elongated feldspars (15%)
with Carlsbad twinning. Sample is less
crystalline than previous sample 10BV-30. Located in the south end of
Horseshoe Basin.
10-BV-32
Trachyandesite taken further up from
previous 10-BV-31. Glassy black
matrix with 15% euhedral plagioclase
up to 5 mm and 5-10% hornblende
laths up to 2 mm. Non-vesicular
jointed outcrop with little weathering.
128
Located in the south end of Horseshoe
Basin.
10-BV-34A
Trachyandesite boulder (20x25cm)
469439E
4459655N
469439E
4459655N
454978E
4457058N
453891E
4458420N
453276E
4459784N
embedded in FCMT welded outcrop.
15% plagioclase feldspar phenocrysts
in a brown-maroon matrix.
10-BV-34B
Trachydacite boulder. Another
volcanic lithic fragment in FCMT,
5cm in diameter, with 15% plagioclase
crystals.
10-BV-39
Basaltic trachyandesite. Maroon
matrix with trachytic texture of
plagioclase euhedral laths (15%),
outcrop is very weathered.
10-BV-40
Basaltic trachyandesite. Fine grained
maroon matrix with 8% euhedral
hornblende phenocrysts (2-5 mm to a
cm megacrysts) and 15% euhedral
plagioclase laths up to 2 mm. Located
in Golconda Canyon, Tobin Range.
10-BV-41
Trachyandesite. Sample is more finely
129
porphyritic and lighter in colour than
previous with a volcanic matrix.
Outcrop is vertically jointed,
weathered, volcaniclastic and
vesicular. Phenocrysts range from 1-4
mm of 5% euhedral hornblende and
10% euhedral plagioclase. Located
further up the Golconda Canyon.
10-BV-42
Basaltic andesite overlain by 33.8 Ma
11449013E
4463684N
11477163E
4424042N
Caetano Tuff. Dark grey vertically
jointed outcrop. Phenocrysts of
plagioclase and olivine weathered to
iddingsite. Black matrix with fine
grained plagioclase, olvine and green
altered mafic minerals. Located at the
west end of the Golconda road.
10-BV-44
Trachydacite taken from a fresh platy
outcrop. Red-dark grey glassy matrix,
nearly aphyric, with 15% euhedral
plagioclase laths (3-5mm), pyroxenes
and possible hornblende up to 1mm.
Minor vesicles. Located off of
130
Highway 305 in the southern
Shoshone Mountains and Red Butte
area.
10-BV-45
Trachyandesite taken from poorly
11474768E
4426305N
474725E
4425612N
474725E
4425612N
exposed, crudely columnar jointed
outcrop. Fine grained, very glassy
black matrix with 15% plagioclase (13mm), trace anhedral olivine <lmm
weathered to brown (5%) and fine
grained pyroxene and olivine. Located
North of Red Butte.
10-BV-46A
Andesite taken from a platy weathered
outcrop. Light grey black matrix with
phenocrysts up to 1 mm of euhedral
plagioclase (15%) and olivine and 1
mm hornblende. Located on the flank
of Red Butte.
10-BV-46B
Trachydacite with more elongated
feldspars than 10-BV-46. Phenocrysts
are not as well preserved in a dark
grey altered matrix. May contain
131
plagioclase megacrysts.
10-BV-47
Rhyolite-trachydacite taken from a
474807E
4425374N
474713E
4424929N
474713E
4424929N
very fissile outcrop. Beds dip west.
Euhedral plagioclase 1-4 mm (15%),
euhedral biotite (10%), some
chloritized, and pyroxene phenocrysts
present in a medium grey crystalline
matrix o f plagioclase and fine grained
hornblende, biotite and pyroxenes.
Dated at 33.82 +/- 0.14 Ma. Located
on a ridge north of Red Butte area.
10-BV-48A
Trachydacite similar to jointed basalts
from a dacitic dome. Contains shear
zones more typical of viscous flows.
Very glassy aphanitic matrix with few
1 mm plagioclase phenocrysts (10%)
with trace biotite, olivine and
pyroxene.
10-BV-48B
Rhyolite taken from a platy
moderately weathered outcrop from
the dome. Contains spherulites, an
indication of divitrification. Also
132
contains phenocrysts of plagioclase,
quartz and biotite in a fine grained
plagioclase groundmass. 28.9 Ma
Campbell Creek Tuff is vertical
adjacent to the dacite with a near
vertical flow attitude making the
dacite a paleo-wall rock. Located at a
lower outcrop of Red Butte.
10-BV-49
Trachyandesite taken from a platy
492020E
4452583N
480460E
4459215N
moderately weathered outcrop. Light
grey matrix with 35% plagioclase, 1-4
mm blocky euhedral plagioclase
phenocrysts, 5% hornblende and trace
l-2mm euhedral pyroxene. Located on
the East side o f Highway 305.
HI 0-53
Rhyolitic Dome with sanidine.
Sampled and dated by Chris Henry.
Dated at 34.24 +/- 0.05 Ma.
133
Appendix B: Thin Section Summary
10-BV-05
Andesite
10-BV-06
10-BV-10
Trachy-dacite
Trachy­
andesite
Trachy­
andesite
Trachy­
andesite
Trachy­
andesite
10-BV-l 1
k-spar, carbonate
lithics
-
5
10
-
3
10
1
2
-
1
1
5
5
20
2
1
5
5
15
1
*■
1
5
15
•
*•
8
2
8
15
-
-
5
Trachy-dacite
-
2
20
-
5
10-BV-12A
Rhyolite
-
-
5
-
10-BV-12B
Andesite
Basaltic
andesite
Trachy­
andesite
Trachy­
andesite
-
2
8
1
5
-
1
-
-
2
10
15
-
■
3
-
1
too fine to tell, glassy and vesicular
1
8
15
1
1
3
-
1
rounded plag
1
8
10
-
•
2
-
1
vesicular, fg plag, qtz, k-spar
vesicular (5%) with glassy vfg
plag
1
10
12
-
•
2
-
1
glassy, vfg plag
lithic clasts
-
1
8
-
1
1
2
1
vfg plag
-
-
10
-
1
2
-
2
.
8
20
.
•
2
-
2
10-BV-07
10-BV-08
10-BV-09
10-BV-l 6
10-BV-29
10-BV-30
-
1
too fg to tell, glassy
•
too fg to tell, glassy and vesicular
2
-
1
vesicular, fg plag
1
fg plag, k-spar, qtz
rounded k-spar
-
1
fg plag, k-spar, qtz
rounded k-spar
5
-
1
vesicular, glassy-vfg plag, k-spar
5
1
2
2
vesicular, glassy, vfg plag
rounded k-spar
rounded k-spar or
plag
1
5
1
1
too fine to tell, glassy
lithic clasts
3
Trachy­
10-BV-31
10-BV-32
10-BV-34A
10-BV-34B
10-BV-39
andesite
Trachy­
andesite
Trachy­
andesite
Trachy-dacite
Basaltic
andesite
fgplagandpyx
134
fg plag, ol
vesicular, glassy vfg felsics
fg felsics of mostly plag,
vesicular
rounded plag
Appendix B continued: Thin Section
Summary
10-BV-40
10-BV-41
10-BV-42
10-BV-44
Basaltic
andesite
Andesite
Basaltic
andesite
_
8
15
3
3
2
.
2
vfg felsics
rounded k-spar and
plag
-
8
15
3
3
2
-
2
vesicular, vfg felsics
carbonate lithics
5
_
10
.
•
•
.
m.g plag and opq
rounded plag
-
5
5
-
-
-
1
vesicular, glassy vfg plag
2
20
.
too fg to tell, glassy
5
15
_
10-BV-45
Trachy-dacite
Trachy­
andesite
10-BV-46A
Andesite
10-BV-46B
Trachy-dacite
-
2
5
-
10-BV-47
Rhyolite
-
2
10
10-BV-48A
Trachy-dacite
-
5
10-BV-48B
Rhyolite
Trachy­
andesite
•
10-BV-49
3
round plag, lithic
clasts
1
1
1
too fg to tell, glassy
1
-
-
1
vfg felsics
-
8
2
1
1
plag and too fg to tell, glassy
2
-
8
2
2
1
plag and too fg to tell, glassy
rounded k-spar
5
5
-
5
2
1
1
plag and too fg to tell, glassy
felsic lithics
10
15
3
3
2
_
1
vfg felsics
Notes: 01 = olivine, pyx = pyroxene, plag = plagioclase, amph = amphibole, bio = biotite, k-spar = potassium feldspar,
qtz = quartz, opq = opaques, - = no phenocrysts were present in thin section
135
Appendix C: X-ray Fluorescence major element data collected in this study
10-BV-05
10-BV-06
10-BV-07
10-BV-08
10-BV-09
10-BV-10
10-BV-ll
10-BV-12A
10-BV-12B
10-BV-16
10-BV-29
10-BV-30
10-BV-31
10-BV-32
10-BV-32
10-BV-34A
10-BV-34B
10-BV-39
10-BV-40
10-BV-41
10-BV-42
10-BV-44
10-BV-45
10-BV-46A
10-BV-46B
10-BV-47
10-BV-48A
10-BV-48B
10-BV-49
H-10-53
00-LT-2+
Avg deviation^
57.1
66.28
58.57
58.58
56.26
55.84
63.29
67.84
54.7
54.89
60.67
60.57
59.28
59.78
59.46
58.19
59.46
52.92
51.93
61.02
52.4
61.23
58.42
56.05
63.67
68.21
67.12
67.7
61.07
70.88
50.95
0.25
0.78
0.62
0.78
0.8
0.88
0.98
0.7
0.42
0.98
1.42
0.84
0.77
0.83
0.83
0.9
1
0.83
1.06
1.04
0.88
1.33
0.6
1.25
0.96
0.56
0.41
0.49
0.41
0.77
0.297
2.41
0.02
13.91
14.97
16.99
17.03
18.71
16.57
15.13
14.76
15.59
15.38
15.21
15.56
15.31
15.2
15.12
18.65
16.93
16.37
17.18
14.43
15.23
15.24
16.31
15.34
16.04
14.99
14.84
14.63
15.27
14.06
13.51
0.07
♦Sample 00-LT-2 is an internal rock standard
♦♦Average deviation for 6 runs o f OO-LT-2
136
6.18
4.71
6.06
6.08
5.71
6.74
4.62
3.67
6.62
8.57
6.47
6.21
6.47
6.65
6.89
5.75
5.91
8.04
8.89
6.69
9.31
5.1
7.74
6.43
4.64
3.26
3.51
2.89
6.1
0.918
13.33
0.05
0.16
0.05
0.1
0.09
0.09
0.13
0.09
0.02
0.11
0.15
0.09
0.08
0.11
0.11
0.12
0.04
0.03
0.14
0.14
0.09
0.15
0.11
0.12
0.12
0.02
0.03
0.05
0.03
0.08
0.01
0.24
0.00
2.07
0.41
2.19
2.01
1.41
2.8
1.83
0.47
3.64
5.08
2.92
2.23
2.98
3.17
3.24
0.55
0.43
3.69
3.79
1.36
4.85
2.15
1.84
3.19
0.7
0.72
1.06
0.66
2.63
0.17
3.94
0.03
Appendix C continued: X-ray Fluorescence major element data collected in this
study
10-BV-05
10-BV-06
10-BV-07
10-BV-08
10-BV-09
10-BV-10
10-BV-l 1
10-BV-12A
10-BV-12B
10-BV-l 6
10-BV-29
10-BV-30
10-BV-31
10-BV-32
10-BV-32
10-BV-34A
10-BV-34B
10-BV-39
10-BV-40
10-BV-41
10-BV-42
10-BV-44
10-BV-45
10-BV-46A
10-BV-46B
10-BV-47
10-BV-48A
10-BV-48B
10-BV-49
H -10-53
00-LT-2*
Avg deviation**
7.32
1.98
5.21
5.37
6.73
6.85
2.38
2.74
5.34
7.84
5.27
4.98
5.36
5.47
5.55
5.2
4.26
6.99
7.24
5.95
8.98
5.13
5.31
6.98
3.68
2.63
3.6
3.02
5.29
1.91
7.67
0.12
2.54
3.49
4.16
4.01
3.5
3.64
4.28
3.53
1
2.89
3.03
3.2
2.99
2.91
2.97
3.24
3.76
3.56
3.52
3.4
3.21
3.45
3.74
1.9
3.58
3.58
3.45
3.42
3.04
3.09
3.24
0.04
2.92
5.25
3.31
3.41
3.33
3.03
4.65
5
4.24
2.19
4.09
4.2
3.87
3.96
3.95
3.65
4.6
2.88
3.09
3.39
0.84
4.3
3.03
3.34
4.74
5.11
4.87
4.99
4.23
4.9
1.92
0.03
137
0.2
0.18
0.49
0.53
0.38
0.51
0.3
0.22
0.39
0.47
0.36
0.36
0.36
0.37
0.39
0.4
0.46
0.62
0.59
0.53
0.24
0.33
0.55
0.37
0.3
0.21
0.25
0.22
0.36
0.09
1.19
0.02
6.81
1.45
1.71
1.11
3.09
2.64
1.43
1.56
7.18
1.17
0.69
0.89
1.78
1.05
0.9
2.17
2
2.56
2.29
1.59
2.85
1.13
1.17
4.61
1.45
1.04
0.98
1.54
0.6
1.69
1.06
0.29
100
99.38
99.57
99.02
100.08
99.71
98.69
100.22
99.79
100.04
99.65
99.05
99.33
99.51
99.5
98.84
98.67
98.83
99.7
99.35
99.39
98.78
99.48
99.29
99.39
100.2
100.21
99.52
99.44
98.01
99.47
Appendix D: ICP-MS trace element data collected in this
study
6.28
1.54
5.24
0.735
4.09
46.35
8.49
1.83
6.69
0.956
5.31
1.02
42.71
7.38
2
5.58
0.718
0.68
44.4
7.7
2.08
5.85
0.758
3.73
3.86
7.08
8.12
1.94
5.83
0.829
4.49
2.19
6.58
0.88
4.68
0.87
0.87
10-BV-05
44.98
83.52
9.43
34.38
10-BV-06
54.36
101.67
12.44
10-BV-07
49.2
94.83
10-BV-08
50.09
97.61
11.3
11.74
0.76
0.71
43.96
40.91
84.7
81.61
10.18
38.68
10.26
41.49
10-BV-ll
60.06
110.78
12.42
43.24
7.24
1.59
5.3
0.707
3.61
0.66
10-BV-12A
45.27
44.54
78.05
86.44
9.19
32.71
5.82
1.34
0.669
3.76
0.78
10.38
39.55
7.31
1.76
4.7
6.03
0.815
4.45
0.82
7.19
0.988
0.96
10-BV-09
10-BV-10
10-BV-12B
13.59
51.88
9.14
2.26
11.15
42.36
7.57
1.76
6.15
0.835
5.26
4.59
10.95
41.88
7.82
1.79
6.35
0.862
4.72
0.9
11.91
44.88
8.23
4.87
40.03
44.05
7.38
8.73
6.52
5.95
0.894
10.87
11.54
1.9
1.72
1.93
7.51
0.815
1.053
4.49
5.57
0.95
0.84
89.02
11.5
44.04
9.44
7.85
6.74
1.055
5.29
0.89
10-BV-16
59.54
10-BV-29
49.28
10-BV-30
47.88
52.32
84.67
94.53
47.89
48.25
91.69
95.07
46.16
10-BV-31
10-BV-32
10-BV-34A
116.26
93.19
0.88
1
10-BV-39
44.12
88.25
44.96
8.49
0.89
4.68
0.87
10-BV-40
41.86
82.68
11.32
10.62
1.83
2.24
42.59
8.28
2.25
6.58
0.871
4.55
0.84
10-BV-41
37.75
76.98
9.17
34.93
6.34
1.67
4.99
0.651
5.3
6.59
5.08
0.763
3.49
4.4
0.64
1.6
1.7
5.3
7.65
0.711
3.86
0.76
1.068
5.97
1.16
6.17
4.46
0.85
0.597
4.7
0.89
4.24
0.565
10-BV-34B
10-BV-42
10-BV-44
27.55
53.91
6.52
43.1
80.48
9.3
10-BV-45
48.11
11.9
8.92
2.25
10-BV-46A
45.32
96.1
87.89
25.53
34.81
46.94
10.49
39.74
7.45
1.78
10-BV-46B
43.48
74.79
8.99
32.21
5.71
1.62
10-BV-47
44.71
10-BV-48A
42.23
78.46
78.12
8.99
31.59
5.47
8.75
31.73
5.69
1.3
1.38
4.47
0.613
1.467
0.01
0.11
0.01
30.56
5.32
1.31
4.08
0.566
10.61
39.65
7.25
1.7
5.69
0.787
H-10-53
55.56
27.06
100.7
10.96
38.28
1.4
8.32
39.1
3.92
3.88
10.24
0.513
58.31
5.91
9.42
0.77
0.05
0.25
0.03
0.08
0.25
138
0.57
0.53
1.76
8.68
88.59
•Sample 00-LT-2 is an inemal rock
standard
••Average deviation for 6 runs of OO-LT-2
0.61
2.82
8.93
79.23
47.09
0.57
0.58
0.56
0.81
43.71
Avg
deviation**
3.15
3.04
3.29
3.08
4.27
10-BV-48B
10-BV-49
OO-LT-2*
0.86
Appendix D continued: ICP-MS trace element data collected in this study
10-BV-05
2.11
0.295
1.918
0.29
90.96
392
13.65
6.32
4.41
1.1
10-BV-06
2.93
0.439
193.96
378
19.51
6.27
8.85
1.3
1.91
0.271
2.849
1.784
0.43
10-BV-07
0.27
1204
12.66
3.45
6.21
10-BV-08
1.98
0.283
1.85
0.28
93.96
102.16
1175
12.8
3.16
6.52
0.7
0.7
5.85
5.43
0.9
1.4
1.1
10-BV-09
2.54
0.358
2.363
0.36
111.18
877
14.42
3.39
10-BV-10
10-BV-ll
2.47
0.347
2.268
0.34
90.58
1100
11.47
3.42
1.85
0.269
212.76
613
17.79
9.95
3.9
2.39
0.378
1.749
2.64
0.26
10-BV-12A
0.43
189.79
559
12.84
6.47
5.31
10-BV-12B
0.306
1.979
0.28
133.52
359
15.02
3.77
4.17
10-BV-16
2.23
2.62
0.357
2.236
0.33
61.85
734
19.27
1.08
10-BV-29
2.47
0.358
2.312
152.02
156.94
660
15.31
6.93
6.81
0.7
1
1
1
10-BV-30
2.55
0.362
2.327
0.35
0.35
661
14.09
5.11
4.55
6.14
1
10-BV-31
0.393
0.352
2.6
0.39
144.59
718
14.98
5.67
6.78
10-BV-32
2.73
2.4
2.302
140.91
663
6.63
2.71
2.23
0.37
2.29
697
3.86
6.89
1
0.299
1.792
0.25
135.31
160.57
14.5
15.34
5.75
10-BV-34A
10-BV-34B
0.35
0.32
1
1
752
14
6.28
0.9
10-BV-39
2.4
0.341
2.225
0.33
71.64
1027
12.23
0.5
6.12
5.71
10-BV-40
2.33
2.006
0.29
0.24
96.02
963
842
10.16
1.99
2.6
10-BV-41
0.311
0.252
12.46
0.38
10-BV-42
1.8
2.4
10-BV-44
2.14
10-BV-45
10-BV-46A
10-BV-46B
1.638
71.85
0.334
2.185
0.32
54.75
514
10.9
4.2
4.63
4.02
2.093
0.32
156.68
6.77
5.58
3.027
0.45
0.342
2.208
0.32
116.18
86.05
760
722
12.12
3.3
2.44
0.307
0.464
1.63
0.225
1.533
0.23
171.96
10-BV-47
1.57
0.228
1.555
0.23
190.2
10-BV-48A
1.77
0.264
1.795
0.27
178.73
10-BV-48B
10-BV-49
1.65
2.34
0.252
1.667
0.26
186.48
552
12.96
0.325
2.188
0.33
153.68
661
14.2
H -10-53
00-LT-2*
Avg
deviation**
1.52
0.227
1.52
0.23
129.98
339
11.74
4.94
0.686
4.393
0.65
42.17
414
7.42
0.02
0.01
0.03
0.01
‘ Samples 00-LT-2 is an internal rock standard
139
0.33
0.7
0.6
0.7
0.7
0.9
1
16.82
2.89
15.28
3.22
6.85
4.94
12.86
6.12
5.77
1
538
13
6.08
5.39
1.1
565
12.7
7.39
5.32
1
7.5
5.38
1.1
4.42
6.15
1
2.25
3.32
0.8
1.73
4.09
0.5
760
751
5.70
0.15
0.03
0.04
1
0.03
Appendix D continued: ICP-MS trace element data collected in this
study
10-BV-05
10-BV-06
2.34
<0.009
0.11
20.5
2.97
0.049
0.2
716.8
1525.3
2.39
0.046
0.07
2.59
0.072
0.16
2.8
0.044
0.13
18.1
20.9
25.33
5.48
28.54
358
21.5
3.14
19.21
266
10-BV-08
11.57
8.9
14.75
20.01
24.24
277
10-BV-09
3.21
3.96
10-BV-10
9.21
2.66
10-BV-l 1
49.98
12.06
24.25
18.89
22.33
12.37
5.91
3.41
236
11
220
16.6
11.8
1561
128
9.4
23.1
24.4
1033.5
1.61
0.081
0.15
2147.1
2.73
0.057
0.17
1605.2
3.13
0.11
0.13
208
22.46
149
295
11.88
2.1
10-BV-29
20.43
5.16
10-BV-30
18.3
20.2
5.16
5.63
19.87
5.37
10-BV-34A
18.1
4.65
26.54
10-BV-34B
16.48
5.82
10-BV-39
9.7
7.92
2.19
22.96
24.44
1.83
22.72
8.74
1.02
17.78
10-BV-42
6.92
1.39
23.12
10-BV-44
10-BV-45
19.51
5.9
11.47
2.96
21.35
31.73
25.66
24.5
25.12
26.46
23.74
1650.9
17.6
14.1
23.92
10-BV-16
10-BV-40
10-BV-41
1571.3
1737.9
164
11.28
10-BV-31
10-BV-32
0.13
20.81
10-BV-07
10-BV-12B
0.043
5.4
9.5
9.1
10-BV-12A
2.55
18.17
8.6
18.2
22
24.6
24
1606.4
2.3
16.9
18.5
2035.2
2.9
0.013
0.028
0.23
272
241
15.1
21
19.4
1595.5
1897
2.65
0.051
0.056
0.21
0.21
270
16.8
261
18.8
1929.7
278
16.6
13.4
2.38
2.26
16.9
1593.5
238
14
73.7
240
16.9
10.9
1567.3
614.8
88
18
9.1
1581
2.07
173
13.2
7.5
1577.6
1889.1
2.29
169
12.9
25
225
13
19.6
1559.9
285
184
16.6
14.1
18.3
9.6
10-BV-46A
13.7
10-BV-46B
21.07
3.69
5.94
24.03
16.64
232
0.1
2.58
0.031
0.091
0.2
0.11
3.52
0.02
0.22
1.37
0.027
0.046
0.23
0.11
0.15
2.81
0.031
0.064
<0.009
0.024
0.21
1404.7
2.31
3.3
15.3
20.4
1824.5
3.05
0.027
0.12
1520.5
3.19
0.07
0.16
0.06
0.16
23
6.08
15.77
208
7.5
22.4
1638.2
2.83
0.087
0.064
10-BV-48A
21.18
6.61
16.84
206
10.5
22
1555.6
2.26
0.035
10-BV-48B
23.63
7.31
16.6
203
8.3
24
1559.4
2.81
0.019
0.12
2.22
0.01
0.12
1.38
0.053
0.18
10-BV-47
10-BV-49
19.37
4.39
22.89
243
15.5
20.3
1338
H-10-53
00-LT-2*
Avg
deviation**
16.17
4.3
4.51
15.76
1388
47.01
4.3
37.6
18.4
1.6
113
154
7
2188.3
0.19
0.06
0.45
2.50
♦Samples 00-LT-2 is an internal rock standard
140
0.40
0.07
96
.
0.1
Appendix D continued: ICP-MS trace element data collected in this study
89
11
21.6
33
15
18.67
19.84
13.2
113
10-BV-08
21.3
18.65
16.54
10-BV-09
15.8
172
65
9
19
11
10-BV-10
55
6
10-BV-ll
7.6
29.4
263
37
19.24
10-BV-12A
16.2
69
10-BV-12B
4.8
48
13
5
19.49
17.77
10-BV-16
10-BV-29
13.6
59
19.87
19.21
2.4
5
0.11
1.54
5985
26
7
5
15.8
5.1
32.1
3.12
16.9
68
11
18.83
11.2
3.15
0.76
0.48
1.88
2.41
3480.6
10-BV-30
3
7
4912.56
10-BV-31
9.9
9.5
48
5
4
21.61
16.4
3.39
4
0.4
21.95
2.31
3
0.21
4843.31
4924.87
5
18.08
22.62
13
33.4
1.83
1.88
2.75
0.52
2.75
2681.07
29.9
1.76
3
5
0.71
7.85
5432.34
10-BV-05
15
10-BV-06
10-BV-07
10-BV-32
10-BV-34A
17.8
2.63
9
0.5
2.35
4677.51
21.6
1.59
6
11.9
3.49
7
0.16
1.31
6306.38
0.51
2.33
3758.21
24
1.07
49
8
1.8
0.42
2.23
2.46
4728.11
5069.52
19.41
22.4
1.99
17.66
31.5
1.61
3
0.46
2.92
2971.89
8.7
11.9
107
7
0.08
1.91
8347.65
3.18
1.74
0.52
2.57
2.8
3
0.53
2.93
5081.5
2610.54
15.8
7.1
29
45
10-BV-34B
8.8
36
9
10-BV-39
39.7
199
50
17.69
7.5
1.17
112
0.09
1.78
8043.49
10-BV-40
12.6
8
21.79
9.5
9
9.4
0.5
2.19
2.35
5859.39
6.7
53
6
18.78
18.54
3
8
7569.1
15.2
3.51
3.24
0.15
10-BV-41
10-BV-42
28
80
33.3
2.49
4
0.37
2.36
3379.66
10-BV-44
10-BV-45
15.8
10.4
58
7
2.03
2.06
6
9
0.57
6513.93
14
38.1
23.5
0.17
62
20.96
20.12
0.7
7.7
59
7
20.6
38.9
2.17
4
0.7
2.72
4.14
4120.85
10-BV-46A
10-BV-46B
5.7
29
4
17.96
36.7
2.67
3
0.33
2.76
2488.78
5643.06
10-BV-47
68
12
18.67
47
7
21.78
10-BV-48A
17
10.2
10-BV-48B
13.8
53
10
19.3
10-BV-49
12.1
28
11
17.71
3
30.4
51
<24
2
4
H-10-53
00-LT-2*
Avg
deviation**
19
21.4
23
22.6
4943.47
7
4
0.6
2.24
0.32
1.84
5169.45
2.61
8
0.34
2.29
4765.01
0.35
4
0.55
6.68
5403.99
0.47
1.83
>12000
3.43
3.11
2
32
-
19.56
9.8
-
-
•Samples 00-LT-2 is an internal rock standard
141
1.39
-
12
2
-
-
-
Appendix D continued: ICP-MS trace element data collected in this study
10-BV-05
0.76
134.41
1.37
83.35
10-BV-06
0.07
0.84
218.25
1.11
137.13
1.86
86.67
1.61
53.49
1.22
91.48
10-BV-08
0.41
106.29
140.64
10-BV-09
0.72
149.29
10-BV-10
0.95
74.55
1.93
69.08
10-BV-ll
0.33
184.61
118.39
10-BV-12A
0.85
148.85
0.61
1.74
10-BV-07
89.68
10-BV-12B
1.06
57.19
2.14
61.53
10-BV-16
165.11
0.9
107.65
10-BV-29
10-BV-30
0.35
0.7
44.2
0.67
147.51
1.89
1.61
78.6
88.84
10-BV-31
0.41
0.74
107.74
10-BV-32
0.49
10-BV-34A
10-BV-34B
1.07
2.24
97.3
95.35
50.34
82.73
10-BV-39
10-BV-40
1.36
0.86
121.8
1.65
82.37
125.96
183.88
2.79
<0.5
0.64
108.48
0.94
10-BV-41
0.52
170.15
1.4
122.59
98.4
10-BV-42
0.91
73.37
1.49
75.58
10-BV-44
0.08
133.41
0.8
95.3
10-BV-45
1.15
89.05
3.41
134.63
80.83
10-BV-46A
3.92
107.43
2.82
110.73
10-BV-46B
1.06
65.85
0.76
48.95
174.64
1.66
10-BV-47
1.5
83.78
10-BV-48A
0.52
99.31
0.99
83.91
10-BV-48B
0.71
134.01
0.23
100.07
1.38
0.54
85.17
10-BV-49
00-LT-2*
0.21
344.42
0.6
Avg deviation**
-
10
-
14
H-10-53
♦Samples 00-LT-2 is an internal rock standard
142
89.07
23
144.62
4
Appendix E: M m Spectrometry ieotepk d a ta caHected ia this study
206/204Pb
2o error
207/704 Pb
2o error
208/204Pb
0.706418
19.401
8.30E-04
15 683
8.24E-04
39.178
2 95E-03
7.29E-04
15.672
8.12E-04
39.039
2 85E-03
143/144Nd
2o error
143/144Nd (i)
87/86Sr
2o error
87/M Sr (i)
lft-BV-06
0.51233859
8.I8E-06
051231406
0.70752387
8.90E-06
0.706807
10-BV47
0.51241803
8.98E-06
0.51239489
0.7063759
1.03E-G5
0.706267
10-BV-08
0.51242100
9.71E-06
0.512397775
0.70638978
8.68E-06
0.706268
I0-BV-10
0.51249190
8.44E-06
0.512465691
0.70604332
1.00E-05
0.705928
10-BV-ll
051238664
8.06E-06
0512364217
0.7069039
1.09E-05
Sampled
2o error
I0-BV-05
10-BV-09
10-BV-12A
051236534
7.36E-05
0.51234)512
0.70705662
9.25E-06
0.706582
19.319
0.70787331
9.01E-06
0.707755
19248
465E-04
15 712
4.83E-04
39359
1 71E-03
19.326
1.14E-03
16.690
120E-03
39091
420E-03
19.361
I.2IE-03
15.709
1.20E-03
39.186
4 14E-03
19.224
5.64E-04
16.678
5.47E-04
38.954
1 94E-03
10-BV-I2B
10-BV-16
051237420
7.92E-06
0512350268
0.706832U
9.51E-06
0.706510
10-BV-30
0.51237877
8.66E-06
0.512353764
0.70690839
9.59E-06
0.706577
10-BV-31
0.51237696
8.67E-06
0512352402
0.70684467
1.00E-05
0.706563
10-BV-32
0.51236361
8.23E-06
0.512338921
0.70681873
1.01E-05
0.706524
10-BV-29
I0-BV-J4A
0.51235794
9 32E-06
0.5123314
0,70709845
9.41E-06
0.706828
IO-BV-34B
051240606
1.01E-05
0.512377355
0.70678282
9 62E-06
0 706484
10-BV-39
051247035
1.08E-05
0512445062
0.7063078
1 53E-05
0.706210
10-BV-40
0.51252269
9.96E-06
0.512496655
0.70590737
126E-05
0705768
10-BV-41
0 51249307
9.01E-06
0.512468763
0.70583258
9.20E-06
0 705713
10-BV-42
051229389
9.52E-06
0512266089
0.70649962
7.53E-06
0.706350
19.397
8.23E-04
15.734
8.26E-04
39471
2 84E-03
10-BV-44
0.51241856
8.91E-06
0512393207
0.70664651
109E-05
0.706358
19.310
9.90E-04
15.682
1.04E-03
39063
365E-03
10-BV-45
0.51238052
1.05E-05
0.512355072
0.70683209
1.24E-05
0.706608
19.317
7.76E-04
15 694
8.27E-04
39.106
2 70E-03
1.16E-03
15.685
1.23E-03
39113
4.16E-03
19.363
10-BV-46A
10-BV-46B
0.51240199
8.67E-06
0.51237825
0.70672991
1.03E-05
0.706410
10-BV-47
0.51238258
9.14E-06
0.512359391
0.70708111
9.22E-06
0.706588
143
Appendix E: M ata Spectrom etry iaetepie d ate t e llected ia this study eon’*
Sampled
143/144Nd
2o error
143/l44Nd (i)
87/86Sr
2o error
87/86Sr(i)
I0-BV-48A
0.51239293
1.02E-05
0.512368915
0.70697525
9.65E-06
0.706533
10-BV-48B
0.51234039
7.90E-06
0.512317077
0.70707142
9.52E-06
0.706601
0.512359133
0.70705737
9.30E-06
0.706732
0512361179
0.707478
10-BV-49
0 51238362
H-10-53
0.512382
Predaioa*
0.511819
9.16E-06
0.710237
0.706951
1 09E-05
* Pb stds = NBS 981 measured over 1 5 yrs, Sr std = NBS 9A7 std measured over 2 yean and Nd std measured over 4 years
144
206/204Pb
2o error
207/204P0
2o error
20S/204Pb
19.361
2.42E-03
16.708
2.71E-03
39.155
2o error
9 73E-03